DEVELOPMENTS IN
QUATERNARY SCIENCES, 6
GLACIOTECTONISM
Developments in Quaternary Science (Series Editor: Jaap J.M. van der Meer) Volumes in this series 1. The Quaternary Period in the United States Edited by A.R. Gillespie, S.C. Porter, B.F. Atwater 0-444-51470-8 (hardbound); 0-444-51471-6 paperback) - 2003 2. Quaternary Glaciations - Extent and Chonology Edited by J. Ehlers, P.L. Gibbard Part I: Europe ISBN 0-444-51462-7 (hardbound)- 2004 Part II North America ISBN 0-444-51592-5 (hardbound)- 2004 Part II1: South America, Asia, Australasia, Antarctica ISBN 0-444-51593-3 ( h b ) - 2004 .
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Ice Age Southern Andes - A Chronicle of Paleoecological Events By C.J. Heusser 0-444-51478-3 (hardbound)- 2003 Spitsbergen Push Moraines -including a translation of K. Gripp: Glaciologische und geologische Ergebnisse der Hamburgischen Spitzbergen-Expedition 1927 Edited by J.J.M. van der Meer 0-444-51544-5 (hardbound) - 2004 Iceland - Modern Processes and Past Environments Edited by C. Caseldine, A. Russell, J. Hardard6ttir, 6. Knudsen 0-444-50652-7 (hardbound) - 2005 Glaciotectonism By J.S. Aber, A. Ber Present volume
Caption for cover image." Aquinnah Cliff on Martha's Vineyard, Massachusetts, United States. The 40-m-high cliff displays multi-colored upper Cretaceous and Tertiary strata that were upthrust along the edge of the Atlantic Coastal Plain during late Wisconsin glacier advance. Aquinnah Cliff is among the most famous and frequently visited glaciotectonic sites in the world. See chapter 6 for more details; photo by J.S. Aber (2005).
Developments in Quaternary Sciences, 6
GLACIOTECTONISM
James S. Aber Earth Science Department, Emporia State University Emporia, Kansas, U.S.A. and
Andrzej Ber Department of Quaternary Geology Polish Geological Institute Warsaw, Poland
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Foreword Between 1987 and 1993 four books with 'glaciotectonics' in the title were published; three were proceedings of meetings, one edited by James Aber, the fourth volume was an edited research volume, the senior editor of which was James Aber. There is no denying that he is an authority on the subject. Since 1993 no general books devoted to the subject have been published. Does that mean that glaciotectonics have gone out of vogue? Certainly not. Checking the reference list of the present volume one will find that numerous papers have been published on all the different aspects of glaciotectonics. In fact there are close to 200 references post-1993 in the list. Not all of these are on glaciotectonics, but neither is it the complete database. These are only the references used to produce the text of the current volume. Glaciotectonics have never been away, they are more important than ever, we were just waiting for someone to pick up the pen and write the next compilation. This has now been done. One shortcoming of all the earlier titles was that relatively little attention was paid to eastern Europe, because there were so few of us in the west who could read the work, if we could get hold of it. By linking up with Andrzej Ber, himself no stranger to studies of glaciotectonics, that extensive eastern European literature has now been made accessible. Comparing the first edition of this volume with the present one will reveal the addition the eastern European literature makes to the subject. The present text will not only be the major reference work on the subject for the years to come, it will also help in defining the next set of questions: What are the major lacunae in our knowledge of the diversity, global distribution and formational processes of glaciotectonics? What are the relations to other glacial landforms, like drumlins, or sediments, like subglacial tills.
Glaciotectonism is the first volume published in Developments in Quaternary Science since my taking over from Jim Rose as Series Editor. Jim gave the series a flying start with the first five titles, one of which consists of three volumes. So far the series is doing very well, with suggestions of reprinting and/or extending some of the first titles. Naturally there is a minor break when changing editors, but now the series is fully on track again. Currently there are seven new titles under contract, and we expect three of these to be published this year. The subjects range from interglacial climates, to early glaciological studies and from China to Patagonia. This range of topics reflects the intention of the series to consider Quaternary science across different parts of the earth with respect to a wide range of Quaternary processes. Nor do we restrict ourselves to the last 2.6 million years. It is obvious that, although that point in time is recognised as a major break, developments since then are not independent of what happened before. Thus, the volume on Patagonia will be on Late Cenozoic developments, not just Quaternary. I hope that the titles published so far and the titles to be published in the near future, will make potential authors consider this book series for their work. Jaap J.M. van der Meer Series Editor
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Contents Preface
ix
Chapter 1: Nature of glaciotectonism Historical development INQUA work groups Definition of glaciotectonism Glaciotectonic terms and concepts Glaciotectonic structures Glaciotectonic landforms Case-history approach
1 1 5 6 7 10 13 15
Chapter 2: Geometric analysis . Introduction Conventional field methods Stereographic projections Subsurface methods Remote sensing Geographic Information Systems
17 17 17 20 22 26 31
Chapter 3: Kinematic analysis . Stress and strain Balanced cross sections Micromorphology Superposed deformation Kineto-stratigraphy Glaciodynamic sequence and event
33 33 35 37 38 40 44
Chapter 4: Hill-hole pair. Introduction Wolfe Lake, Alberta Herschel Island, Yukon Devils Lake, North Dakota Norwegian continental shelf
45 45 48 50 53 55
Chapter 5: Composite ridges . . Introduction MCns Klint, Denmark Dirt Hills and Cactus Hills, Saskatchewan Flade Klit, Denmark Utrecht Ridge, Netherlands Brandon Hills, Manitoba
59 59 63 67 72 75 80
Chapter 6: Cupola hills and drumlins Introduction Aquinnah, Martha's Vineyard, Massachusetts Ristinge Klint, Denmark Hvideklint, Men, Denmark Elblgg Upland, Poland Saadj~irve drumlin field, Estonia
83 83 83 88 89 93 99
viii
Aber and Ber
Chapter 7: Megablocks and rafts . Introduction Esterhazy, Saskatchewan Southern Alberta Kvarnby, Skhne, Sweden Sinim~ied, Estonia Chapter 8: Intrusions, diapirs and wedges . Introduction Atchison, Kansas Herdla Moraine, Norway Systofte, Falster, Denmark Kronowo esker, Poland
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Chapter 12: Dynamism of Glaciotectonic Deformation Fundamental cause of glaciotectonism Initiation of thrust faulting Continuation of thrust faulting Scale models of glaciotectonism
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Chapter 13: Glaciotectonic analogs Introduction Mississippi Delta mudlumps, Louisiana Thin-skinned thrusting Convergent plate boundary References .
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141 141 142 145 147
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Chapter 11: Distribution of glaciotectonism Continental distribution of glaciotectonic phenomena Regional patterns of glaciotectonism Model for lobate pattern of glaciotectonism Glaciotectonic patterns in North America Glaciotectonic patterns in central Europe Central Europe m Weichselian Glaciation Central Europe - - Saalian and Elsterian glaciations Glaciotectonic patterns in Arctic Russia
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125 125 125 127 130 133
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Chapter 10: Applied glaciotectonics Introduction Highwall failure, Highvale coal mine, Alberta Highway construction, Maymont, Saskatchewan Diatomite quarries, Fur, Denmark
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111 111 114 117 120 121
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Chapter 9: Basement and deep crustal structures . Introduction Canadian Shield and northern Appalachians Northern Scandinavia Salt displacement, Finger Lakes region, New York Polish basement structures
Index
101 101 102 104 107 109
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153 153 160 162 167 168 171 181 187 191 191 193 196 198
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203 203 203 205 210
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2 1 5
235
Preface
This book is dedicated to our glaciotectonic mentors, Profs. Asger Berthelsen, University of Copenhagen, Denmark and Hanna Ruszczyfiska-Szenajch, University of Warsaw, Poland (figs. 1 and 2). During the 1970s and '80s they led a renaissance in glaciotectonic investigations and theory in the Baltic region of northern Europe (fig. 3). Their contributions and enthusiasm for the subject were major influences for the authors and many other geoscientists. Many other colleagues were involved with our work on glaciotectonics. In particular, we wish to acknowledge the contributions and support of the following individuals. Inge Aarseth (Norway) Ojars P. Aboltins (Estonia) David L. Ackerman (Canada) Harold V. Andersen (U.S.A.) L~erke T. Andersen (Denmark) Valery Astakhov (Russia) Peter H. Banham (United Kingdom)
Figure 1. Asger Berthelsen on the beach near his country home on the island of MOn, Denmark, where he conducted much of his glaciotectonic research. Photo by J.S. Aber (1979).
John E Bluemle (U.S.A.) Julie Brigham-Grette (U.S.A.) Krzysztof Brodzikowski* (Poland) Bruce E. Broster (Canada) Chris D. Clark (United Kingdom) David G. Croot (United Kingdom) Louis F. Dellwig (U.S.A.) Linda A. Dredge (Canada) Alexis Dreimanis (Canada) Mark M. Fenton (Canada) Darek Gat~zka (Poland) Jane K. Hart (United Kingdom) Stephen R. Hicock (Canada) Howard C. Hobbs (U.S.A.) Michael Houmark-Nielsen (Denmark) Ole Humlum (Denmark) Mads Huuse (Denmark) William R. Jacobson, Jr. (U.S.A.) Peter Roll Jakobsen (Denmark) Wojciech Jaroszewski* (Poland) Carrie Jennings (U.S.A.)
Figure 2. Hanna Ruszczyfiska-Szenajch examines a freshly dug exposure at Kadyny in Elbl~g Upland, one of her many field sites for glaciotectonic investigations in Poland. Photo by J.S. Aber (1993).
x
JCrn Bo Jensen (Denmark) Flemming JCrgensen (Denmark) Volli Kalm (Estonia) Alexander Karabanov (Byelrussia) John S. Klasner (U.S.A.) Rudy W. Klassen (Canada) Alan R. Knaeble (U.S.A.) Stefan Kozarski* (Poland) Johannes Krtiger (Denmark) Dariusz Krzyszkowski (Poland) Erik Lagerlund (Sweden) E.A. Levkov* (Belarus) Tom van Loon (Netherlands) Ida LCnne (Norway) Kam Lulla (U.S.A.) Jan Lundqvist (Sweden) Holger Lykke-Andersen (Denmark) Jan Mangerud (Norway) Leszek Marks (Poland) Irene Marzolff (Germany) Andrei V. Matoshko (Ukraine) Jaap van der Meer (United Kingdom) Rod A. McGinn (Canada) David M. Mickelson (U.S.A.) Wojciech Morawski (Poland) Daniel Nyvelt (Czech Republic) Stig A. Schack Pedersen Jan Piotrowski (Denmark) Mikko Punkari (Finland) Maris Rattas (Estonia) Bertil Ringberg (Sweden) Karol Rotnicki (Poland) Joar Saettem (Norway) David J. Sauchyn (Canada)
Aber and Ber
Nicholaus M. Short (U.S.A.) Steen SjCrring* (Denmark) Everett E. Spellman (U.S.A.) Hans-Jtirgen Stephan (Germany) James T. Teller (Canada) William A. Thomas (U.S.A.) F.M. (Dick) van der Wateren (Netherlands) William A. White (U.S.A.) Richard S. Williams, Jr. (U.S.A.) Vitalijs Zelcs (Latvia) * deceased. Financial support for our efforts has come from many sources, not the least our home institutions, Emporia State University, Kansas (Aber) and the Polish Geological Institute, Warsaw (Ber). A series of grants from NASA provided Aber with experience and tools for remote sensing and GIS applications in the geosciences. Aber's experience in Estonia was funded by the U.S. National Research Council, and his field work in Poland was supported by the U.S. Council for Intemational Exchange of Scholars. INQUA provided material aid to Ber in connection with the glaciotectonic mapping project in central Europe. Other sources of support include the North Dakota Geological Survey, Danish Natural Science Research Council, and William T. Kemper Foundation - Commerce Bank, Trustee - Kansas City, Missouri. Thanks to the Swiss Geological Society for permission to reproduce copyrighted illustrations. Finally we wish to thank our wives, Susie and Maria, who patiently assisted us during preparation of this book and encouraged our long-term commitment to understanding glaciation and its role in shaping the Earth.
Figure 3. Strongly sheared and refolded chalk-till m~lange exposed in cliff near Hvideklint on the island of Men, southeastern Denmark. Small spade to right for scale. Photo by J.S. Aber (1979).
Chapter 1 Nature of Glaciotectonism Historical development Glaciotectonism is an important component of modern glacial theory, but it gained widespread recognition only within the past 25 years. Glacial theory began to develop in the late 18th and early 19th centuries in the Alps of western Europe and the mountains of southern Scandinavia. Horace-Bdn6dict de Saussure was among the earliest naturalists to undertake systematic observations of glaciers in the Mont Blanc vicinity beginning in the 1760s. He observed types of glaciers, ice flow, origin of moraines, and many other aspects of glaciers. Saussure introduced the terms roches moutonn6e, s6rac, and moraine into geological usage (Carozzi and Newman 1995). James Hutton was first to recognize in 1795 that erratic granite boulders in the Jura Mountains had been transported by glaciers from the Alps (Flint 1971). In Scandinavia, Jens Esmark concluded in 1824 that glaciers once had been much larger and thicker, and had covered much of Norway and the adjacent sea floor (Andersen 1992). He attributed erratics and moraines to glacial transportation and deposition. Esmark also recognized that glaciers were powerful agents of erosion that had carved out the Norwegian fjords (Cunningham 1990). In the early 1830s, Jean de Charpentier began to marshall the scientific evidence for former alpine glaciation. He observed moraines, striations, and erratics, as well as existing glaciers of the Swiss Alps, and presented his conclusions in a persuasive manner (Teller 1983). Although a disbeliever at first, Louis Agassiz came to accept the concept of former alpine glaciation, and then carried the idea much further. In 1837 he proposed that vast sheets of ice had once covered much of the northern hemisphere. This was a radical suggestion, for at that time the modem ice sheets in Greenland and Antarctica were completely unknown. Agassiz undertook detailed studies of glacier movement on the Unteraar Glacier in Switzerland in the 1840s, and he influenced James D. Forbes to begin similar glaciologic research in the French Alps. Forbes established that glaciers move in part by internal viscous (plastic) deformation, in contrast to the more popular dilatation or regelation theories of the day (Cunningham 1990). The glacial theories of Esmark, Charpentier, Agassiz and Forbes were based on recognition of three features--large erratic boulders, moraines, and abrasion marks on boulders and bedrock, interpreted in light of observations on glacier dynamics. Taken together these features could be explained only as the results of formerly more-extensive glaciation.
Glacial theory from the beginning, thus, rested on two groups of geological field evidence: 1) features formed by glacial erosion and 2) features created by glacial deposition. The possibility that glaciers could deform shallow crustal rocks and sediments was not recognized until a few decades later. Charles Lyell (1863) was among the first geologists to discuss the origin of contorted glacial strata at Norfolk, England, in the Italian Alps, and elsewhere. He suggested three possible mechanisms for deformation: 1) pushing by stranding icebergs, 2) melting of buried ice masses, and most importantly 3) pushing before advancing glacier ice. Lyell's The Antiquity of Man (fig. 1-1) was widely read, and his comments on glacially deformed sediments must have influenced many geologists. During the late 19th century, ice-pushed structures were recognized in several now classic locations" Sk~ne, southernmost Sweden (Torell 1872, 1873; Erdmann 1873); MCns Klint, Denmark (fig. 1-2) and Rtigen, Germany (Johnstrup 1874); and southern New England islands in the United States (Merrill 1886a). The obvious structural disturbances at these places previously had been ascribed to a variety of causes, including landslide, volcanism or intrusion, mountain building, etc. Recognition of ice-shoved features from interior continental locations came about in the latest 19th and earliest 20th centuries. Gilbert (1899) noted glacier-deformed bedrock structures on the southern shore of Lake Ontario, New York. Various ice-pushed structures were described near Minneapolis, Minnesota (Sardeson 1905, 1906) and in Poland (Frech 1901, 1915). Hopkin's (1923) analysis of large iceshoved hills in eastern Alberta is perhaps one of the best early studies. He emphasized the similarity in structural style of these hills compared to the foothills of the Canadian Rocky Mountains. The first geologist who really specialized in the study of glacially deformed structures was George Slater, an Englishman. He chose the subject of ice-push deformation for his doctoral dissertation at the University of London. He studied ice-shoved hills in England (Slater 1927a, 1927b), Denmark (Slater 1927c, 1927d), Canada (Slater 1927e), the United States (Slater 1929), and the Isle of Man (fig. 1-3; Slater 1931). He was the first to use the term glacial tectonics (Slater 1926), which is now generally shortened to glaciotectonics (American) or glacitectonics (British).
2
Aber and Ber by observations of neoglacial push-moraines on Spitsbergen made by the German geologist Gripp (1929). However, when Gry moved to a position at the Geological Survey, teaching of glaciotectonics ceased at the University, and the study of Danish glaciotectonics languished for many years thereafter.
OF
THE ANTIQUITY OF MAN WITR KKKhREB ON T K l O R I B 8 0 l p
THE ORIGIN OF SPECIES BY VARIATION
BY SIR CHARLES LYELL, F.R.S. .LUTKOIt 0Y ' P ~ I K C I P L E 8 0 l t GKOLOG~," ' KLKMKNTS 0 1 GEOLOGY," ITC. ITC.
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Figure 1-1. Cover page of Lyell's "The Antiquity of Man" (1863), which contains an early discussion of contorted drift.
Unfortunately, the Danish geologist Axel Jessen, who had once helped Slater, later accused Slater in no uncertain terms of plagiarizing and misrepresenting his (Jessen's) work at LCnstrup Klint, Denmark (fig. 1-4; Jessen 1931). In spite of Slater's prolific publications, glaciotectonism was still considered an unusual or peculiar manifestation of the glacial theory. Small contortions in glacial strata were commonly recognized, but many geologists continued to deny that glaciation could create large deformations or build substantial ice-pushed landforms. The growth of glaciotectonic research in Europe during the middle portion of the 20th century was checkered by distinct national differences. The study of glaciotectonic phenomena emerged briefly as a specific field of research in Denmark during the 1930s under the leadership of Helge Gry (Univ. Copenhagen) and Jessen (Geological Survey Denmark). Gry's (1940) analysis of ice-shoved hills in the Limfjord region of northwestern Denmark demonstrated the full potential of combining structural geology with glacial stratigraphy and geomorphology. Gry was strongly influenced
Before World War II, glaciotectonic research in what is now Poland was carried out by German geologists (Bergrat and Illner 1928; Woldstedt 1932; Meister 1935; Berger 1937; Schwarzbach 1942) as well as Polish geologists (Lewifiski and R6~ycki 1929; Czajka 1931; Rr~ycki 1937). Glaciotectonics next emerged as an important subject of research in the Netherlands following World War II (Crommelin and Maarleveld 1949; de Jong 1952; Maarleveld 1953). The reason for Dutch interest in glaciotectonics is obvious: ice-shoved ridges are the most conspicuous topographic features in a country otherwise noted for its flatness. The Netherlands continues to be a center for glaciotectonic research, as shown by numerous detailed investigations (e.g. Ruegg and Zandstra 1981; van der Meer 1987; van der Wateren 1992), and Dutch scientists have played a leading role in Svalbard glaciotectonics (van der Meer 2004). Glaciotectonics continued as a significant field of research in Poland (Jahn 1956; Dylik 1961; Galon 1961; Krygowski 1965; Rotnicki 1976) during this period. Glaciotectonics is among the most active research fields in Polish glacial geology today (Ruszczyfiska-Szenajch 1979; 1985; Ber 1987, 1999; Jaroszewski 1991, 1994; Krzyszkowski 1996). Icepushed structures are increasingly important for overall Quaternary geology (Brodzikowski and van Loon 1985, 1991; van Loon and Brodzikowski 1994). Meanwhile, research in northern Germany reflected a renewed interest in ice-push deformation there (e.g. Hannemann 1970; Grube and Vollmer 1985; Stephan 1985; van der Meer 1987), and similar efforts were underway in Belarus of the Soviet Union (e.g. Lavrushin 1971; Levkov 1980; Karabanov 1987). Glaciotectonic research in Denmark was revived during the 1970s by Asger Berthelsen (Univ. Copenhagen), who applied his experience with hard-rock structural geology to unraveling glaciotectonic phenomena. He developed the method of kineto-stratigraphy (Berthelsen 1973, 1978), in which the main emphasis is placed on the study of the directional elements that reflect the movement patterns (kinetics) of former ice sheets (1978, p. 25). Berthelsen motivated students, and his kineto-stratigraphic method was highly successful for working out Weichselian glacial stratigraphy in Denmark. In fact, glaciotectonic analysis is now an integral part of geological mapping in Denmark (Petersen 1978) and plays an important role for Quaternary studies (Houmark-Nielsen 1988, 1994, 1999; Pedersen 2000).
Nature of glaciotectonism
3
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Figure 1-2. Earliest known illustration of MOns Klint, from Pontoppidan's "Danske Atlas" (1764). Chalk mass in center, Sommerspiret (B), stands in a vertical position with the pinnacle >100 m above sea level.
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Figure 1-3. Structural sections through the Bride Moraine at Shellag Point, Isle of Man, United Kingdom. Above - section from Slater (1931). Below - section measured by Thomas (1984). Note similarities in these sections described more than 50 years apart. Adapted from Thomas (1984, fig. 3). Reproduced from Boreas by permssion of Taylor & Francis AS.
4
Aber and Ber
is approximately twice that of combined Fennoscandian/ British Ice Sheet coverage (Flint 1971). Beginning in the 1950s, large-scale topographic maps and aerial photographs became widely available, which greatly facilitated geologic reconnaissance. At the same time, many states and provinces commenced systematic geologic mapping programs. The North American glaciotectonic renaissance began in western Canada on two fronts. Working in southern Saskatchewan, A.R. Byers (Univ. Saskatchewan) demonstrated the ice-thrust genesis of the Dirt Hills and Cactus Hills. These are among the largest and best-developed glaciotectonic hills in the world (Byers 1959). Walter Kupsch (Univ. Saskatchewan) expanded on Byers' work to include similar large ice-shoved hills across southern Saskatchewan and eastern Alberta (Kupsch 1962). At the Saskatchewan Research Council, reconnaissance mapping, done mainly by Earl Christiansen, resulted in much additional knowledge of glaciotectonic features (Christiansen 1961, 1971a, 1971b; Parizek 1964; Christiansen and Whitaker 1976).
Figure 1-4. LOnstrup Klint, North Sea coast of northwestern Denmark. This cliff erodes rapidly, retreating at a rate > 1 m per year. A - northward overview of cliff section which exceeds 80 m height and extends several km. B - detail of tilted Quaternary strata exposed in the cliffface near Rubjerg Knude; north toward left. A complete glaciotectonic analysis of this classic cliff section was completed recently by Pedersen (2005). Photos by J.S. Aber (2004).
Following Slater's controversial career, British interest in glaciotectonics was minimal. The classic chalk rafts and contorted drift along the Norfolk coast received some attention in the 1960s (Peake and Hancock 1961; Harland, Herod and Krinsley 1966), and a revival in glaciotectonic studies there was led by Peter H. Banham (Univ. London). Banham (1975, 1977) also applied hard-rock geologic principles to interpretation of glaciotectonic structures. Banham's work complemented that of Berthelsen along with the Soviet geologist Lavrushin (1971). Together, they built the methodologic foundation for modem glaciotectonic research. Aside from several early studies on southern New England islands (Hollick 1894; Woodworth 1897; Upham 1899)and a few other isolated investigations, little glaciotectonic research was carried out in North America until the late 1950s. Part of the explanation may be the fact that fewer glacial scientists were faced with a much larger and less developed geographic area to investigate. The glaciated region covered by the Laurentide Ice Sheet in North America
At the same time, equally impressive ice-shoved ridges on the Yukon Coastal Plain were described by J.R. Mackay (1959). These features, developed wholly within permafrost, were the basis for Mackay and colleague W.H. Mathews (both Univ. British Columbia) to develop a general theory for glaciotectonic thrusting (Mathews and Mackay 1960; Mackay and Mathews 1964). They were influenced strongly by Hubbert and Rubey's (1959) analysis of overthrust faulting in mountains. The Saskatchewan discoveries soon spilled over into North Dakota, where a county mapping program was underway. Led by Lee Clayton (Univ. North Dakota), Steven R. Moran and John L. Bluemle (both North Dakota Geological Survey), the state became a productive center for glaciotectonic research during the 1960s and '70s. Many local studies were conducted, culminating in a new Geologic Map of North Dakota (Clayton, Moran and B luemle 1980), on which a variety of glaciotectonic landforms is shown. Theoretical analysis accompanied the field mapping (Moran 1971; Clayton and Moran 1974; Bluemle and Clayton 1984). When Moran moved to the Alberta Research Council, the geographic area was expanded to include the entire glaciated Great Plains region of the United States and Canada. The result was the first attempt at continent-scale synthesis concerning ice-sheet dynamics, distribution of glacial landforms, and genesis of glaciotectonic phenomena (Moran et al. 1980). Clayton later moved to the Wisconsin Geological Survey, with the unsurprising result that glaciotectonic features are now recognized in that state. With increasing surface and subsurface information, it now appears that glaciotectonic phenomena are ubiquitous in glaciated regions underlain by sedimentary bedrock or thick
Nature of glaciotectonism
"'"~
DEFORMATION
5
+1
Figure 1-5. Triad of effects created by glaciation, on which modern glacial theory is based.
drift (Moran 1971; Moran et al. 1980) and are even common in thin drift resting on crystalline bedrock. A variety of distinctive landforms is now attributed either wholly or partly to glaciotectonic genesis. Hence, glaciotectonic features must be included with depositional and erosional features as primary field evidence for glaciation. The modern glacial theory is, therefore, supported by a triad of field evidence, including erosional, deformational, and depositional features (fig. 1-5). Although hardly recognized in older glacial geology textbooks (e.g. Flint 1971; Drewry 1986), glaciotectonism has achieved a prominent place since the 1990s (e.g. Hambrey 1994; Benn and Evans 1998; Evans 2003).
INQUA work groups International recognition of the significance of glaciotectonic phenomena came in 1982, when a Work Group on Glacial Tectonics (WGGT) was established within the International Union for Quaternary Research (INQUA). Berthelsen organized WGGT and served as its first President. The overall goals of WGGT were: to initiate and stimulate studies of glaciotectonic phenomena in both recent and ancient glacial environments, to promote interdisciplinary collaboration between scientists working in different parts of the field, and to increase glaciotectonic curriculum in academic teaching and professional training (WGGT Newsletter, 3/1987). Leadership of the work group was taken over by David G. Croot (Plymouth Polytechnic, United Kingdom) in 1987, who was succeeded in turn by Jane K. Hart (Univ. Southampton, United Kingdom) in 1995. WGGT's formal existence came to an end in 1999, when its activities were incorporated into other work groups of the INQUA Commission on Glaciation.
Figure 1-6. WGGT work group field conference at MCns Klint, Denmark. During a lunch break, A. Berthelsen (standing to right) points out a feature in the cliff exposure. Photo by J.S. Aber (1986).
During its existence, many activities and publications were sponsored under the WGGT umbrella (fig. 1-6). Special symposia and field conferences were conducted (e.g. Sjcrring 1985; van der Meer 1987; Croot 1988a; Aber 1993a; Sauchyn 1993; Warren and Croot 1994), and glaciotectonic examples have been used for laboratory exercises in structural geology (Aber 1988a). The importance of glaciotectonic structures and landforms is now well accepted among many researchers actively engaged with studies of glacial geology and geomorphology. One of the original goals of WGGT was to assemble a comprehensive bibliography of references to published reports, books, maps and other documents pertaining to glaciotectonics. This effort was pursued through a series of publications (Aber 1988c, 1993b), and with the emergence of Internet the glaciotectonic bibliography moved online in the late 1990s [http://www.geospectra.net/]. Another major goal of WGGT was to assemble regional and continental geographic databases on glaciotectonic phenomena. This effort was initiated at a work group meeting on the island of Men, Denmark in 1986 and formally adopted at the INQUA Congress in Canada (1987). Project organization began immediately in North America. At that time, a controversial decision was taken to pursue development of the map and database using Geographic Information System (GIS) methodology. In the late 1980s, GIS still was considered experimental and risky for application to such a large and complex undertaking. David J. Sauchyn and David L. Ackerman led construction of the GIS database at the University of Regina (Canada).
6 Primary compilers for glaciotectonic data were B luemle, Lynda Dredge (Geological Survey Canada), Julie BrighamGrette (Univ. Massachusetts) and Aber. They adopted a morphological approach for classifying and mapping glaciotectonic phenomena (Aber 1988b). In many cases, glaciotectonic features in North America are known mainly or only from their morphologic expression, so structural classification schemes proved ineffective.
Aber and Ber of Stig A. Schack Pedersen (Geological Survey Denmark) and then by Croot (Croot and Michalak 1993). It became apparent that compiling a European database on glaciotectonic phenomena would be more complicated than in North America. Agreement on how to proceed was more difficult to achieve among various countries with different scientific traditions in mapping glacial geology and geomorphology. Diverse points of view and working methods hampered initial progress.
An initial map and database for the northern Great Plains region was completed in 1991 (Aber et al. 1991). This first effort provided a template upon which to elaborate the whole glaciated region of North America. A preliminary dataset for the whole continent was achieved two years later (Aber et al. 1993), and a final map was published just in time for the INQUA Congress in Berlin, Germany (Aber et al. 1995). All the GIS datasets were placed online and remain available for anonymous FTP from the Data Access and Support Center (DASC) at the Kansas Geological Survey [http:// gisdasc.kgs.ku.edu/].
The demise of the Soviet empire in the early 1990s was a foremost event for integrating geological research in western, central, and eastern Europe. The emergence of newly freed countries and their scientific establishments gave renewed impetus for the glaciotectonic mapping project. Russian geologists collaborated with scientists from the west for glaciotectonic investigations, for example Valery Astakhov and Jan Mangerud (Norway). Similar transnational cooperations developed between many other European and North American countries.
Success of the GIS approach for the North American glaciotectonic map led to creation of a new work group at the Berlin congress. A work group on Geospatial Analysis of Glaciated Environments (GAGE) was approved as part of the INQUA Commission on Glaciation. GAGE was organized by Aber, who became its first president, and later he was joined by Ber as co-president. This work group continued the goal of assembling GIS datasets and producing maps of glaciotectonic structures and landforms (fig. 1-7). GAGE remained quite active until the INQUA Congress in the United States (2003), when GAGE officially completed its mission.
At the INQUA Congress in Berlin in 1995, Ber assumed the role as main coordinator of the glaciotectonic mapping project for central Europe with assistance from many individuals. The resulting glaciotectonic database spans the Baltic region from Denmark to Estonia and southward to the Ukraine and Czech Republic. Glaciotectonic maps have been prepared for individual countries (Jakobsen 2003; Ber 2003a; Ber and Krzyszkowski 2004; Rattas and Kalm 2004), and a first version of the glaciotectonic map of central Europe was presented in 2003 at the INQUA Congress in the United States (Ber 2003b; Ber and Aber 2003).
Meanwhile, organization of the European glaciotectonic mapping project began to develop, first under the leadership
Definition of glaciotectonism Glaciotectonism refers to the processes of glaciotectonic deformation. However, considerable uncertainty surrounds the exact meaning and use of the term glaciotectonic, and confusion is both semantic and conceptual. Semantically it is both an adjective and noun. In English, the word glaciotectonic is an adjective that should be used to modify a noun, for example glaciotectonic landform. In noun forms, glaciotectonics refers to features or results of glaciotectonic deformation, whereas glaciotectonism refers to the processes of glaciotectonic deformation.
Figure 1-7. GAGE work group field conference, Slovakia. Group prepares to take a raft tour down the Dunajec River in the Tatra Mountains. Juraj Janocko (standing to right) was the conference organizer. Photo by J.S. Aber (2000).
Conceptual confusion arises because various deformed structures are common both in glacier ice and in glacial deposits. Although Slater (1926) did not define the phrase glacial tectonics, he used it in reference to structural disturbances in both drift and bedrock as well as deformations in glacier ice. Similar, if somewhat vague, references have been given by other geologists in the years since. An important review by Occhietti (1973) included five categories of glacially related deformation.
Nature of glaciotectonism •
• • •
•
- deformation of pre-existing substratum (drift and bedrock) by active ice movement. glaciodynamic- primary structures (such as till fabric) produced within ground moraine by active ice. glaciostatic - deformation of ground moraine and substratum by static ice loading. g l a c i o k a r s t i c - deformation accompanying freezing and thawing of buried dead-ice masses. i c e b e r g d r i f t i n g - deformation of sea or lake sediments by grounded icebergs. glaciotectonic
The latter two categories may be eliminated from further consideration, leaving threeAglaciotectonic, glaciodynamic, and glaciostatic, which are the primary topics of this book. is an all-encompassing term defined here as glacially induced structural deformation of bedrock or sediment masses as a direct result of glacier-ice movement or loading (based on Moran 1971; Aber, Croot and Fenton 1989). Maltman, Hubbard and Hambrey (2000) adopted a quite similar definition with the further stipulations that structures within glacier ice are not included and primary depositonal structures (such as till fabric) are excluded. Glaciotectonics
The word "bedrock," as used in this definition, implies any type of pre-existing consolidated or unconsolidated materials ranging from crystalline basement rock, to lithified sedimentary rock, to unlithified or loose sediment; in other words, the substratum and foreland over which the glacier advances. Glaciotectonic deformation also may involve sediments deposited during the same glaciation that subsequently deformed those strata. A glacier or ice-sheet may induce deformations of substratum or foreland m~terial as a result of its forward (dynamic) movement or its vertical (static) loading. Both causes of deformation often operate simultaneously, and the effects of each cannot usually be separated. Therefore, glaciodynamic and glaciostatic structures are considered as joint manifestations of secondary deformations produced during glaciation. Depression and rebound of the crust and lithosphere, as a consequence of glaciation and deglaciation, are forms of glaciotectonism, long recognized and much studied (Walcott 1969; Andrews 1970; Tushingham 1992). However, the emphasis in this book is placed on shallow (200 160120 100 eo-
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Figure 2-17. Seismic section, TEM soundings, and test drilling data from a profile across the buried Viuf valley in eastern Jylland, Denmark. A steeply dipping reflection (near 1200 m in profile) was interpreted as an inclined raft of Paleogene clay based on discrepancies in the levels of Paleogene clay found in the test drilling and in surrounding TEM soundings. Taken from JCrgensen et al. (2003, fig. 5); reprinted with permission from Elsevier. 1300
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Figure 2-18. Seismic section across the FanO Bugt thrust complex, offshore southwestern Denmark. Two ddcollement levels are evident. The basal ddcollement dips uniformly toward the west; the shallow ddcollement exhibits considerable relief Numerous inclined thrusts and pop-up structures are identified. Depth in two-way-travel (TWT) time; 400 ms is ~360 m depth. Taken from Andersen (2004, p. 45).
scope of this text to describe the plethora of remote sensing techniques that might be applied to glaciotectonic situations. For further details about remote sensing, see Jensen (2000) and Lillesand, Kiefer and Chipman (2004). Following are brief descriptions of some common methods.
Aerial photography has been a mainstay in geological exploration and mapping programs for more than half a century. In single frames or stereopairs, conventional airphotos are essential for all types of modern geological and geomorphological investigations, including study of glacial
Aber and Ber
28
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Figure 2-19. Seismic profile from the Ortowo vicinity, southern Warmia, Poland. Q - Quaternary, N - Neogene (Miocene), E Paleogene (Oligocene, Eocene, Paleocene), K - Cretaceous. Note slices of Oligocene strata (E~) thrust and folded within Miocene (N) toward right side of profile. Adapted from Morawski (2004b, fig. 10). -
Figure 2-20. Conventional panchromatic airphoto of the Crestwynd vicinity, southern Saskatchewan, Canada. Large composite ridges cross the scene from NW to SE. Photograph A21639-7 (1970); original photo scale = 1:80,000. Reprocessed from the collection of the National Air Photo Library, Natural Resources Canada.
Geometric analysis
29
Figure 2-21. Landsat multispectral scanner (MSS) scenes from southern Saskatchewan. A - spring image, 21 May 1978, MSS band 7 (near-infrared). The Dirt Hills and Cactus Hills are large ice-shoved ridges on the northeastern margin of the Missouri Coteau upland. B - winter image, 9 Dec. 1972, MSS band 7 (near-infrared). Note the "pen-and-ink" quality of this snowcover view in which topography is quite evident. This is among the earliest Landsat images. Scanned from original image prints at scale 1:1,000,000; acquired from U.S. Geological Survey, EROS Data Center.
30
Aber and Ber
Figure 2-22. Manned space-shuttle photograph of southern Denmark and northernmost Germany. Low-oblique view toward north; compare with Fig. 2-6. Late Weischelain ice lobes followed Baltic Sea channels between the islands and mainland. Derived from color-visible photograph, STS68-153-038, Sept. 1994. Image obtained from NASA Johnson Space Flight Center, courtesy of K. Lulla.
landforms (Way and Everett 1997). This mature method has expanded in recent decades upward to space and downward to low-level platforms in order to provide diverse kinds of imagery. Conventional airphotos are medium-scale, panchromatic, large-format (23 cm), vertical views (fig. 2-20). Glaciotectonic landforms have been pictured in airphotos for geologic atlases, for example Gravenor, Green and Godfrey (1960) and Prest (1983). Such airphotos are indispensible for regional mapping and assessment for all manner of environmental conditionsmsoils, geology, water resources, vegetation, etc. Aerial photographs of this type are available in principle for most northern countries in which glaciotectonic landforms are conspicuous; however, the age, quality and cost of such airphotos vary greatly. At a nominal scale of 1:40,000, for example, a vertical airphoto covers approximately 85 km 2 ground area, and 0.1 mm on the airphoto represents 4 m on the ground. The usual resolution limit for conventional airphotos at scales 1:20,000 to 1:25,000 is 1-2 m; smaller objects cannot be discerned
unless they have high contrast with the surroundings. The panchromatic (black-and-white) nature of conventional airphotos also limits their use for interpreting and classifying vegetation cover, which is often an indicator for underlying geologic conditions. Color-visible or color-infrared aerial photographs are available for selected areas in only a few countries. A recent trend is conversion of hardcopy airphotos into digital orthophotos, which have been scanned and rectified to fit a cartographic grid system. Orthophotos may be imported and used directly in GIS databases (see below).
Space-based remote sensing for civilian, scientific applications began with the manned Skylab Missions and the unmanned Landsat satellite series in the early 1970s (fig. 2-21). Landsat instruments provide a wealth of moderateresolution datasets for all manner of earth-science applications. Many Landsat datasets are in the public domain and are available from the U.S. Geological Survey EROS Data Center or from other remote sensing centers at modest cost. The potential of Landsat imagery for glacial geomorphology was recognized early by Morrison (1976) and Slaney (1981). In regions of low relief, winter Landsat images
Geometric analysis
31
impart a strong morphologic aspect to the scene due to low sun angle, snow cover, and lack of active vegetation (Skoye and Eyton 1992). Manned space photographs likewise have given regional overviews for visualizing glacial landscapes (fig. 2-22). Other satellite systems of particular interest are Radarsat, operated by the Canadian Centre for Remote Sensing, and Ikonos, a commerical, high-resolution multispectral scanner. The spatial resolution of the latter rivals traditional airphotos. Many other systems and types of space-based remotely sensed imagery are widely available nowadays from various governmental and commercial sources. Remote sensing from space provides several advantages compared to conventional aerial photography (Short and Blair 1986). • Synoptic, big-picture views that portray large areas of the Earth's surface. • Multispectral data--visible, infrared and microwave portions of the spectrum. • Repetitive, global coverage throughout the year and spanning several decades. • Low cost compared to collecting same type and volume of data on the ground. The application of space-based remote sensing has revolutionized geomorphology, and indeed has led to a new subdiscipline known as megageomorphology, which is concerned with geomorphic description and analysis of large areas (Baker 1986). This approach has extended to glacial geomorphology (Williams 1986), in which remote sensing has provided the data for large-region to continent-scale mapping and interpreting of glaciated landscapes. Punkari (1980) was among the earliest to use this approach for interpreting glacial dynamics in Finland, which he subsequently elaborated to include all of Scandinavia and northwestern Russia (Punkari 1996). Similar investigations were carried out in North America (e.g. Boulton and Clark 1990a, 1990b; Clark 1993). At the other end of the height and scale spectrum, smallformat aerial photography (SFAP) has emerged as a means to acquire large-scale, high-resolution imagery for detailed site investigations (Warner, Graham and Reed 1996; Bauer et al. 1997). This method is based on 35- or 70-mm film cameras as well as compact digital or video cameras operated from low-height, manned or unmanned platforms. Many types of lifting platforms may be employed, such as manned aircraft (Light 2001), and unmanned model airplanes (Quilter and Anderson 2000), blimps (Ries and Marzolff 2003) or kites (Aber et al. 1999). Vertical photographs typically have ground resolution of 2-10 cm (linear pixel size), which is suitable for mapping and analysis in the microstructural scale range, 1:100 to 1:1000 (Masing 1998).
Figure 2-23. Kite aerial photograph of Feggeklit, a small ice-shoved hill on the island of Mors, northwestern Denmark. The cliff section on the eastern (right) side exposes deformed Eocene Fur Formation uplifted from the floor of the Limfjord estuary in the background. For more information see chapter 3. Photo by S.W. Aber (2005). Relatively low cost combined with high portability and ease of use in the field make SFAP an excellent method to acquire large-scale airphotos for site-specific geomorphic investigations (Marzolff and Ries 1997). Glacial landforms, such as eskers and drumlins, are portrayed especially well in SFAP (Aber and Gat~tzka 2000; Aber and Kalm 2001), and the method is well suited for depicting glaciotectonic landforms (fig. 2-23). SFAP has proved effective for accurately mapping complex glaciotectonic structures exposed in high cliff sections. The Danish approach involves low-height, oblique airphotos taken in overlapping sequence along the cliff face. The photographs are acquired from a small, manned airplane that flies parallel to the cliff at a distance of 200-300 m. Ground control points are derived from standard vertical airphotos of the study site. Multi-model photogrammetric analysis is applied to oblique stereopairs based on the method developed by Dueholm (1992). This technique was utilized for construction of detailed geological sections for two of Denmark's most famous glaciotectonic sitesnMCns Klint (Pedersen 2000) and LCnstrup Klint (see fig. 1-4). In the latter case, Pedersen (2005) produced a section 6 km long at scale 1:500 with a spatial resolution of ~25 cm. Detailed geometric data of this type are invaluable for elaborating the glaciotectonic genesis of such key sites.
Aber and Ber
32 An important approach for remote sensing is the multiconcept, including multilevel, multispectral, and multitemporal datasets (Avery and Berlin 1992). Each type of remote sensing has particular strengths for application to glacial geomorphology, and used in combination these systems provide data spanning all spatial, spectral and temporal dimensions of observation.
.................................
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Geographic information systems The multiconcept may be expanded to include many other types of relevant data, such as field observations, subsurface logs, topographic maps, soil maps, digital elevation models, etc. Multiple kinds of evidence can be integrated in geographic information systems (GIS), which allow display, analysis, and modeling of complex landscapes based on many "layers" of input data. GIS has emerged rapidly during the past two decades to become a powerful tool of choice throughout academic, governmental and commercial enterprises dealing with earth science data and applications (Burrough 1986; Faust, Anderson and Star 1991; DeMers 2003; Clarke 2004). The potential for utilizing digital elevation models (DEM) was demonstrated by Lidmar-Bergstr6m, Elvhage and Ringberg (1991) for display and interpretation of glacial landforms in Skhne, southern Sweden. At the time, their DEM was a classified military database, but now moderate- to highresolution DEM datasets are available to the public for many portions of the world. Aber (1999) employed a shaded-relief map derived from a DEM as a background for portraying preIllinoian glacial landforms in the central United States. Colgan, Mickelson and Cutler (2003) expanded the GIS approach for the southern Laurentide ice sheet from the Dakotas to New England. Clark (1998) realized the potent capability of GIS for reconstructing and modeling paleo-ice sheet dynamics over continent-sized areas. GIS was adopted for compilation of glaciotectonic maps in North America (Aber et al. 1991, 1995) and in central Europe (Ber and Abet 2003). As an example, Jakobsen (1996) employed a GIS approach for mapping the density of concealed (subsurface) glaciotectonic structures in Denmark based on well records (fig. 2-24). The distribution of concealed glaciotectonic structures reveals some distinctive patterns. In many cases, areas with high densities of concealed structures coincide with well-known glaciotectonic landforms (various ice-shoved hills). But in some other cases, zones with intense
~>
Figure 3-12. Model for rotational strain within a bed undergoing differential movement (faster at top, slower at bottom). The rotations may involve individual grains or particle aggregates. Adapted from van der Meer (1997, fig. 2); reprinted with permission from Elsevier.
Micromorphology
Figure 3-10. Photomicrographs of highly strained diamicton in cross-polarized light. Above - sigmoidal grain (s) suffered internal strain. Arrows indicate sense of shear. Field of view ~2½ mm across. Middle - galaxy structure. Note spiral fabric (concentric layers) of grains around the core clast (c). Field of view ~4 mm across. Below- reworked diamicton pebbles (d) embedded within plasma. Field of view ~4 mm across. Images courtsey of W.R. Jacobson, Jr.
Microscopic structures within deformed glaciogenic materials have proven valuable for discerning the manner of strain imposed by glacial stress, and a large body of research has been conducted on this subject (e.g. Menzies and Woodward 1993; van der Meer 1993, 1997; Menzies 2000). The method involves collecting oriented samples, which then are impregnated with epoxy and cut into thin sections. The thin sections are examined with a petrographic microscope to describe and measure the physical and structural properties of the deformed material (Khatwa and Tulaczyk 2001). On this basis, a large number of characteristic brittle and ductile microstructures are recognized (fig. 3-9).
stacked thrust faults dipping in the upvalley direction. Total shortening is 54%. This example demonstrates the power of balanced sections to evaluate kinematics in glaciotectonic settings.
These studies have demonstrated that much strain takes place by rotational movements within the deformed material. Rotations may be internal to a single grain, take place between adjacent particles, or involve particle aggregates. Typical
Aber and Ber
38 microstructures include sigmoidal grains, galaxy fabric, and reworked clasts (fig. 3-10). Shearing may produce pressure shadows (fig. 3-11). Differential movement within a deforming layer may be likened to a series of turning wheels or ball beatings (fig. 3-12). The styles and orientations of such microstructures may be utilized to determine how the material has strained, and this could lead to reconstruction of the stress regime.
Superimposed deformation One of the great challenges in structural geology is working out multiple phases of deformation, each of which may have a different character and orientation. In shields and mountains each phase of deformation imparts a new set of structures that overprint or even obilterate older structures. Glaciotectonic structures display similar possibilities. Pedersen (2000) elaborated four settings for superimposed glaciotectonic structures (fig. 3-13). • Glaciotectonic structures superimposed on preglacial tectonic structures. • Glaciotectonic structures representing two distinct regional ice advances.
• Glaciotectonic structures for different phases of deformation by the same ice advance, • Glaciotectonic structures superimposed by postglacial neotectonic structures. In addition to these, glaciotectonic and neotectonic structures could develop simultaneously in some situations. In all these settings, multiple phases of deformation lead to complicated structural patterns that can be quite difficult to tease apart from field observations based on limited exposures, test drilling, GPR, or other techniques. In fact, excellent exposures are often necessary to unravel such details, for example in long cliff sections or large quarry faces. Among the most remarkable structures are refolded folds, that is folds affected by two (or more) phases of deformation from different directions. A dramatic example of this superimposed structure is the "arrowhead" fold pattern (fig. 3-14). Crossing fold sets create interference patterns within a rock or sediment body, just like crossing wave sets on a water body. Erosion of the folds may reveal such interference as arrowhead or crescent patterns (fig. 3-15). Refolded structures may, in rare cases, be evident in the landscape, where fold crests coincide with hill crests, as en 6chelon patterns (Pedersen 2000).
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Figure 3-13. Four settings for superimposed glaciotectonic structures. A- glaciotectonic structures superimposed on preglacial tectonic structures, B - glaciotectonic structures representing two distinct ice advances, C- glaciotectonic structures representing different phases of deformation within the same ice advance, D - glaciotectonic structures superimposed by postglacial neotectonic structures. Adapted from Pedersen (2000, fig. 1).
Kinematic analysis
39
8 2
Figure 3-14. Remarkable exposure of an "arrowhead" fold pattern in Eocene strata, island of Mors, Denmark. Oblique photograph showing a horizontally cleaned surface in a claypit. See next figure for geometric explanation. Image courtesy of S.A.S. Pedersen.
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Figure 3-15. Model for refolded structure. Fold phase 1 consists of tight, overturned folds with gently dipping axial planes, which are warped by fold phase 2, an upright, open anticline. Erosion across the top of the resulting folds reveals the interference pattern as arrowhead and crescent outcrop patterns. Compare with photograph above. Taken from Pedersen (2000, fig. 7). Pedersen (1993, 1996) demonstrated the challenge and reward of deciphering superimposed deformation at Feggeklit, island of Mors, Denmark. A sequence of deformed Paleogene and glaciogenic strata are well-exposed in a coastal cliff section, nearly 1 km long (see fig. 2-23). Dislocated bedrock consists of the Fur Formation comprised of distinctively interlayered clayey diatomite, tephra beds, and cemented zones. Numerous tephra beds provide stratigraphic markers throughout the
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Figure 3-16. Stratigraphy of Fur Formation and glaciogenic strata exposed at Feggeklit, northwestern Denmark. 1 southern portion of cliff section, 2 - middle section, 3 northern section. Individual tephra beds are numbered for stratigraphic correlation. Adapted from Pedersen (1993, fig. 3).
formation. Olst and Holmehus Formations underlie the section. The glaciogenic sequence includes three tills" T1 Saalian, dislocated; T2 - probably Saalian, dislocated; T3 Weichselian, discordant (fig. 3-16). The cliff section trends NNE-SSW parallel to the direction of glacier movement, and thus normal to strike/trend of most structures. Several phases of deformation took place as Weichselian ice advanced from the NNE and eventually overrode the vicinity. Initial phases of deformation generated various fractures, folds and faults, and took place in a proglacial environment. The final phase was subglacial and produced breccia (glaciotectonite). On this basis, Pedersen
Aber and Ber
40
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Figure 3-17. Photograph of deformed Fur Formation in southern portion of Feggeklit section, zone 1. Dark tephra layers are interbedded with light-colored diatomite. Note the low amplitude, long wavelength folds and small faults. Height of cliff exposure ~10 m. Photograph by J.S. Aber (in Pedersen 1993, plate 13).
ANASTOMOSING J~NT!I~
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(1996) recognized three structural zones, from least disturbed in the distal (southern) part of the section (zone 1), to moderate deformation in the central portion (zone 2), to greatest dislocation at the proximal (northern) end of the section (zone 3). Each section is represented by characteristic structures.
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• Zone 1 - - long wavelength, low amplitude folds, anastomosing joints, and conjugate faults (fig. 3-17). • Zone 2 - thrust fault ramp, conjugate faults and box folds, listric splay thrust faults. • Zone 3 - - low-angle thrusts and splay thrust faults, overturned anticlines.
Figure 3-18. Fracture and fault structures at Feggeklit (below) showing typical dip angles (0). Mohr circles (above) showing relationship between angle of failure (20) and increasing ice pressure (Cru). Adapted from Pedersen (1993, fig. 7).
The dip angles (0) of fractures and faults may be related to increasing ice pressure during advance toward and over the vicinity (fig. 3-18). Balanced cross sections for Feggeklit are the basis for constructing a basal d6collement, which is projected to lie in bentonite of the upper Holmehus Formation, ~100 m deep (fig. 3-19). The balanced sections indicate approximately 100 m of horizontal shortening, which represents about 10% overall strain for the Feggeklit structural complex. Based on varves and other chronologic control, Pedersen (1996) estimated that deformation took place within a 50-year time span and that regional ice advance averaged 10 m per year, a reasonable value for steady-state (non-surging) ice movement.
stratigraphy in a routine manner. Individual till sheets cannot be traced far. Morphologic features may be relicts of older glaciations. A glacial advance that created much deformation may leave few deposits of its own. In many ways, the problems of stratigraphy in glaciotectonic terrain are analogous to stratigraphy of complexly deformed crystalline shields. The stratigraphy of shields is often considered as a sequence of deformational events. Each event is marked by a distinctive style, type and orientation of structures and is characterized by certain metamorphic or igneous rocks.
Kineto-stratigraphy In many regions with abundant glaciotectonic phenomena, it may be difficult or even impossible to conduct glacial
Berthelsen (1973, 1978) developed a similar stratigraphic approach, called k i n e t o - s t r a t i g r a p h y , for unraveling complex Quaternary sequences in glaciotectonic terrain. He defined a kineto-stratigraphic drift unit as, the sedimentary
unit deposited by an ice sheet or stream possessing a characteristic pattern and direction of movement (Berthelsen 1973, p. 23).
Kinematic analysis
41
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Figure 3-19. Schematic balanced cross sections for Feggeklit, northwestern Denmark. The pin point is placed at the southern end of the exposed cliff sequence. 1 - simplified stratigraphy and structure of the exposed section. 2 - balanced cross section with thrust faults projected to a basal ddcollement. Also shown are antithetic (backthrust) faults. 3 -final deformed section. The final section is approximately 100 m shorter than the balanced section. Adapted from Pedersen (1996, fig. 17).
A kineto-stratigraphic drift unit includes all primary sediments (till + stratified drift) that were deposited during all stages of a particular glacier advance and retreat (fig. 320). The primary sediments may be deformed by the same ice advance subsequent to deposition. Such deformations are termed domainal; deformation of subjacent and older strata is extra-domainal. The lower boundary of domainal sediments is the limit of penetrative deformation, which also marks the bottom of the kineto-stratigraphic drift unit (compare with fig. 1-9). The unifying factor in kineto-stratigraphy is the directional character of the deposits that can be related to ice movements
during a particular glaciation. The directional character is revealed by many kinds of features: striations and grooves, till fabrics, indicator erratics, and most importantly glaciotectonic structures. Domainal deformations are most useful in this regard, as there is no doubt about the glacier advance to which they correspond, but extra-domainal structures may be used also for kineto-stratigraphy. The kineto-stratigraphic principle of Berthelsen (1978, p. 29) is that: Deposits laid down by successive glaciations can be distinguished by means of the kinetic patterns deduced from the domainal and extra-domainal deformation
\ S
A, 3
Az '
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Figure 3-20. Schematic cross section of deposits classified in a kineto-stratigraphic drift unit. Black = till; stipled = stratified drift. A -- advance phase; S = maximum (sandur) phase; R = recessional phases; K = kame. Ice advance from right to left. Taken from Berthelsen (1978, fig. 1).
Aber and Ber
42
related to each glaciation, and that the glacial deposits should be correlated according to their directional elements.
opposite flanks may differby as much as 120 °. Near the lobe axis, ice flow direction will not change much during the course of lobe advance and retreat. However, flow directions will shift significantly on lobe flanks as the lobe expands and shrinks. These local patterns of ice movement and glaciotectonic deformation are consistent in light of the lobate model for glaciation (see chap. 11).
In order for kineto-stratigraphy to be applied successfully, three prerequisites must be met. First, consecutive ice advances must come from different directions. Multiple advances from the same direction would all be classified in a single kineto-stratigraphic drift unit. Second, each advance must display a consistent and recognizable regional pattern of movement. And third, it must be possible to relate various directional features to the proper ice advance. This is a matter for careful field observations and structural analysis.
Since its inception, kineto-stratigraphy has been applied for working out glacial advances in diverse situations. Hicock and Dreimanis (1985) demonstrated the advantages of this approach in British Columbia and Ontario, Canada. In fact, they found glaciotectonic structures are more reliable than till fabrics for establishing the direction of ice movement. Glaciotectonic structures may be measured quickly in the field and serve as a check on other directional indicators. Elsewhere in North America, middle Pleistocene ice-lobe advances were deciphered with kineto-stratigraphy in the central Great Plains of the United States (Aber 1991).
The second p r e r e q u i s i t e ~ c o n s i s t e n t pattern of ice movement~may be the most difficult to deal with, because kineto-stratigraphy is an empirical method. It is tested regionally, but developed by fitting together local results. The local results may at first seem inconsistent or even haphazard when considered in isolation. The direction of ice movement may vary markedly across a region and may even vary significantly at individual sites for different phases of the same glaciation.
Recognition of glaciotectonic structures proved invaluable for unravelling glacial stratigraphy and glacier dynamics in the Kap Ekholm sections, Spitsbergen, Svalbard, where a sequence of marine and glacial strata are exposed in cliff sections up to 30 m high and more than one km in total length. This key section had been studied before in considerable detail, but the presence of glaciotectonic structures was not recognized by earlier investigators (Boulton
Consider an idealized ice lobe advancing from the north. A northerly flow direction is developed only along the central axis of the lobe. In the lobe flanks, ice flow diverges toward the southeast and southwest. Thus, flow directions on
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Figure 3-21. Stratigraphy and structures exposed in a portion of the Kap Ekholm sections, Spitsbergen, Svalbard. Low-angle thrusts and small folds are located within Formation F, especially in the 20-22 m level. The sandy diamicton (shaded) at the top of the formation is interpreted as a deformation till. Formation F consists of mid-Weichselian marine strata that were deformed by late Weichselian glacier overriding. Adapted from Mangerud and Svendsen (1992, fig. 15); reprinted with permission from Elsevier.
43
Kinematic analysis
1979), which led to some erroneous stratigraphic interpretations. Mangerud and Svendsen (1992) identified low-angle thrusts and normal faults as well as various folds and vertically tilted strata. Many of these structures had escaped previous recognition, because they strike/trend N-S, parallel to the main direction of cliff exposures (fig. 3-21). The orientation of these structures demonstrated late Weichselian ice movement from east to west, crosswise to the adjacent fjord, which suggests a relatively thick ice cap that moved independently of local topography. Kineto-stratigraphy has proven most successful in the country of its origin--Denmark (Berthelsen 1978; Houmark-Nielsen 1981, 1987). Five phases of Weichselian glaciation are recognized (fig. 3-22). Old Baltic ice lobes moved into the southern islands from the southeast. Two old Baltic advances are now recognized and considered to be mid-Weichselian in age (Houmark-Nielsen 1994, 1999). Late Weichselian glaciation took place in several stages beginning with the Norwegian advance coming from the north and covering northern Sj~elland, Sams¢, and Jylland (Pedersen 2005).
with at least one or two readvances and a shift to more easterly movement during the course of ice coverage. Following a brief interstade, the Young Baltic ice lobe overspread the islands from the southeast and reached eastern Jylland. The final Ba31thav readvance took place in the form of ice tongues moving from the south along channels between larger islands and the mainland. All these late Weichselian advances happened during a relatively short time interval between 20,000 and 14,000 years BE The kineto-stratigraphy of Denmark is based on the traditional model for Weichselian glaciation (Holmstr6m 1904; Andersen 1966), in which ice lobes were fed by ice streams following topographic depressions and emanating from the interior of the ice sheet. This model is best illustrated by the Young Baltic phase (fig. 3-23). The ice lobe advanced from east to west along the southwestern Baltic, and divergent ice flow turned toward the northwest and north in southern Denmark. Local variations in ice m o v e m e n t and glaciotectonic deformation could be explained by development of small ice tongues on the margin of the main lobe.
The Main Weichselian advance next came from the northeast and reached into central Jylland. This advance was complex
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Aber and Ber
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Figure 4-17. The "Horseshoe" at Trcenabanken. Sea floor is approximately 300 m deep in vicinicty of the ice-shoved hill which has a strongly curved outline. The presumed source depression (hole) is located on the partly enclosed, concave side of the hill. Taken from Scettem (1990, fig. 2). The "Horseshoe" is a hill with strongly curved outline on the eastern portion of Tr~enabanken (fig. 4-17). The middle, northwestern portion of the hill stands more than 100 m above the surrounding sea floor, which is about 300 m below sea level. The northern margin of the hill is a steep escarpment that extends eastward into a low ridge; the southern margin curves to the southeast also into a lower ridge. This arrangement is strongly arcuate in outline, and the concave side partly encloses a topographic basin. The hill covers an area of approximately 20 km 2, has a volume of ~1 km 3, and rests on a substratum composed of Quaternary sediment. A shallow sediment core from the hill crest revealed till, which had moderate undrained shear strength (10-90 kN/m 2) The basin associated with the Horseshoe occupies an area about 35 km 2 and has an estimated volume of 0.7 km 3. Glaciomarine or marine clay is sparse on the hill and within the basin. The Horseshoe and the adjacent sea-floor depression represent a well-preserved hill-hole pair of nearly ideal form. The steep distal (northwestern) margin of the hill and its abrupt boundary with undisturbed sea floor imply the hill was not overridden by the glacier margin subsequent to deformation.
The main portion of the hill is, thus, interpreted as a result of pushing at the front of the last glacier advance to reach the vicinity. This advance took the form of an ice tongue. The lower limbs of the hill may have formed by lateral squeezing along the flanks of an ice tongue or as material was scraped aside during advance of the ice-tongue front (see fig. 4-8). The moderate undrained shear strength of hill sediment is compatible with fast ice flow, perhaps a surge, and rapid sediment deformation. On FuglCybanken, Steinbitryggen is an irregular ridge that rises about 90 m above the adjacent sea floor (fig. 4-18). The ridge extends approximately 30 km east-west, covers an area of about 58 km 2, and has a volume of about 1.6 km 3. At the eastern end of Steinbitryggen is a depression, known as Sopphola, that covers 56 km 2 in area and has a volume of about 1.7 km 3. The main mass of Steinbitryggen has a triangular or arrowhead shape, includes several distinct ridges, and is located adjacent to Sopphola. Sopphola comprises several discrete basins. The arrangement of Sopphola and the proximal mass of Steinbitryggen suggest ice pushing from southeast to northwest. However, the point of the arrowhead is elongated into a low ridge that extends
Hill-hole pairs
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~~i~.',' " ~ ,:.~ ......... , -7:.{..., .100 m high. The pinancle of Sommerspiret fell down during a storm in the early 1990s. B - Grcederen chalk mass with pinancle 82 m high. Photos by J.S. Aber (1986). 20,000 to 13,000 years BP (SjCrring 1981). Hcje Men generally exceeds 75 m elevation reaching a high at 143 m (fig. 5-9). Dronningestolen (the Queen's throne), the largest chalk cliff in the center of MCns Klint, is 128 m high (fig. 510). Presumably undisturbed chalk bedrock is situated 20 to 40 m below sea level (Haarsted 1956; Jensen 1993), so >160 m of structural relief is indicated.
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The individual scales exposed in MCns Klint consist of upper Cretaceous (Maastrichtian) white "writing chalk" that was deformed along with drift. Chalk now forms ridges and cliffs because of its greater resistance to erosion, whereas intervening drift has been eroded into valleys that form the falls along the coast (fig. 5-11). The chalk is quite uniform in lithology, aside from occasional layers of flint nodules, and thus, stratigraphic correlation between chalk masses is
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Figure 5-8. MCns Klint section as viewed from the east; Jcettebrink at southern end (upper left), Slotsgavle at northern end (lower right). Black lines within blank chalk masses show deformed flint layers. Reproduction of copper engraving by Puggaard (1851); adapted from International Geological Congress XXI Session, Norden, Guidebook I (1960).
65
Composite ridges
difficult at best. Surlyk (1971) divided the Danish Maastrichtian into ten brachiopod biozones: zones 1-7 are lower Maastrichtian and 8-10 are upper Maastrichtian. These biozones are the best method for establishing correlation between chalk masses.
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,'100 m above sea level. In relation to adjacent source basins, total relief exceeds 200 m in some cases (van der Meer 1988-89). Among these impressive glaciotectonic landforms is Utrecht Ridge, located on the southwestern edge of Gelderse Vallei in the central Netherlands (fig. 5-29). It is one of a series of ridges composed of imbricately thrust scales of unconsolidated Pleistocene strata that are situated around three sides of a glacial basin located in Gelderse Vallei (Ruegg 1991). Utrecht Ridge trends northwest-southeast, and it rises >50 m above the adjacent lowlands, forming a conspicuous sight in an otherwise flat landscape. Utrecht Ridge was created when an ice lobe excavated Gelderse Vallei basin. The ice-shoved hills are heavily eroded as a consequence of their Saalian age combined with unconsolidated character of Quaternary strata and are mere ruins of the original landforms. Kwintelooijen sand pit is located on the inner or northeastern side of Utrecht Ridge
Aber and Ber
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Figure 5-27. Block diagram depicting the 3-dimensional extent of the Hanklit thrust complex based on cliff exposure (front face) and geo-electric survey (toward back). Note the slight undulation of the Hanklit thrust sheet, which rises approximately 20 m toward the eastern end. Takenfrom Klint and Pedersen (1995, fig. 16); copyright Geological Survey of Denmark and Greenland.
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Composite ridges
77
Figure 5-28 (right). Buried valleys on the island of Mors, northwestern Denmark, identified on the basis of large-scale TEM mapping. H-S = Hanklit-SalgjerhOj glaciotectonic complex. Asterisk (*) shows location of young buried valley immediately west of H-S. Adapted from JOrgensen et al.
Valley generations Youngest generation Oldest generation
(2005, fig. 9).
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(fig. 5-30). The pit came under coordinated, interdisciplinary study during the 1970s and early 1980s on account of the many Paleolithic artifacts and fossils found there (Ruegg and Zandstra 1981).
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The sedimentary strata now forming Utrecht Ridge were originally deposited as alluvium of the ancestral Rhine/Meuse Rivers. Kwintelooijen sand pit contains three formations, in ascending order: Kedichem, Urk, and Drente (Ruegg 1981). The Kedichem Formation consists of very fine sand, clay, peat, and loam deposited in a flood-plain/back-swamp environment. Sediments of the Urk Formation are mainly fine to coarse sand, gravelly sand, and coarse gravel deposited by the ancestral Rhine River (Zandstra 1981). The glaciofluvial Drente Formation contains fine to coarse sand, gravelly sand, and gravel. These formations total >20 m in thickness.
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The Drente Formation, by definition, includes all drift related to the Saalian glaciation. At Kwintelooijen, it accumulated as ice-marginal and sandur deposits in front of glacier ice.
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Figure 5-29. Ice-pushed hills and glacial basins of the central Netherlands. K = location of Kwintelooijen sand pit within Utrecht Ridge on the southwestern side of Gelderse Vallei. Based on van den Berg and Beets (1987).
Aber and Ber
78
features demonstrate that Utrecht Ridge probably was not overridden by the Gelderse Vallei ice lobe (van der Wateren 1985), although post-Saalian erosion may have removed an original till cap. Utrecht Ridge consists of thrust blocks striking parallel to the ridge and dipping on average 35 ° to 40 ° NNE (fig. 531). Thickness of scales varies from about 25 m to only a few m, but each includes a basal portion of fine-grained sediment of the Kedichem Formation. Thrust blocks are imbricately stacked and gently folded. Thrusts at the base of each scale contain shear planes, small isoclinal folds, breccia and slickensides in a zone of intermingled sediments several dm thick. Many normal faults forming conjugate sets are also present.
Figure 5-30. Topographic map of Utrecht Ridge showing location of Kwintelooijen sand pit (K) and position of section (fig. 5-31). Elevations in m; contourinterval = 5 m. Modified from van der Wateren (1981, fig. 1).
Much of the Drente sediment was probably reworked from underlying Urk Formation deposits. In southern Gelderse Vallei, Saalian till is buried 15-30 m below sea level. A test boring made at the bottom of the sand pit penetrated thrust and contorted strata to a depth of about 24 m below sea level (Zandstra 1981). Utrecht Ridge near Kwintelooijen is about 2.5 km wide with a plateau top between 45 m and 60 m above sea level (fig. 530). The ridge is a s y m m e t r i c in cross profile. The northeastern side slopes 5-15 ° , whereas the southwestern flank slopes only 2-5 ° . This difference is partly explained by the presence of sandur deposits covering the southwestern portion of the ridge (van der Wateren 1981). Utrecht Ridge and other ridges surrounding Gelderse Vallei are cut by several dry valleys representing former spillways, and the ridge plateaus show no trace of till. The morphologic
The present elevation of Utrecht Ridge is less than its initial elevation, as a result of lowering by normal faulting and by later erosion. Van der Wateren (1985) estimated that scales were initially pushed up at least 100 m above the basal drcollement. Structural restoration results in a good balance between volumes of Utrecht Ridge and the excavated basin of southern Gelderse Vallei. Van der Wateren (1981) calculated that potential shear stress developed at the glacier sole would be far too small to upthrust blocks 100 m above the drcollement. Instead of shear stress, the lateral pressure gradient caused by differential loading of the substratum during ice advance provided the driving force for glacier thrusting. A series of deep glacial basins behind ice-shoved hills stretches across the central Netherlands (fig. 5-29). The basins and associated hills increase in size from west to east. Owing to greater subsidence, ice-pushed ridges of the western basins are now mostly buried beneath younger sediments. These basins exceed 100 m depth and are floored with Saalian till in many places. The basins also contain tunnel valleys that appear to lead toward spillways breaching the marginal ice-shoved hills.
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Composite ridges
79
Prior to the Saalian glaciation, the Netherlands was an alluvial plain with slight topographic relief, having no major valleys or hills. The northern Netherlands was underlain by mostly fine-grained sediments over which the Saalian ice sheet advanced easily. As the ice sheet reached the central Netherlands, it overran coarser, gravelly sediment (Urk Formation) of the ancestral Rhine/Meuse Rivers. This caused an increase in basal friction, thickening of the ice sheet, and melt-water erosion of subglacial basins and tunnel valleys. The marginal effect of this modified subglacial topography was to generate ice lobes in a region where the landscape was previously flat. The thicker ice sheet in combination with melt-water erosion of subglacial basins and development of ice lobes was an ideal situation for thrusting of composite ridges above a basal d6collement in the Kedichem Formation. Two mechanisms contributed to thrusting: 1) lateral pressure gradient due to differential loading (van der Wateren 1985), and 2) direct glacier pushing against the sides of the basins (van den Berg and Beets 1987). Thus, the glacial basins were created by joint subglacial melt-water erosion and glaciotectonic thrusting. Certain factors remain unresolved, for example the nature of permafrost. Most Dutch geologists have assumed that deformation took place in permafrozen sediments and that the thickness of scales is an indication of the depth of permafrost (de Jong 1967). Van der Wateren (1985), however, challenged the assumption of permafrost as a prerequisite for ice thrusting.
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Figure 5-32. Map of the Netherlands showing locations of ice-shoved hills and directions of ice movement. Five phases of ice-shoving are noted from south to north (a-e). Adapted from Aber (1985, fig. 2).
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More recent models, such as van den Berg and Beets (1987), de Gans, de Groot and Zwaan (1987), and Rappol et al. (1989) include only two or three glacial phases by ice streams coming across the North Sea basin or crossing from Germany. Laban
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Jelgersma and Breeuwer (1975) explained the origin of glacial basins as deep erosion caused by glacier surging, and this is hardly compatible with ice advance over permafrost. Another possibility is that the ice sheet was cold-based during the advancing stage. Later, when the ice warmed and became more mobile, glacier lobes cut deep basins and pushed up ridges (ter Wee 1983). Five phases of Saalian ice pushing were recognized previously in the Netherlands (fig. 5-32), three in the central portion and two in the northeastern region (Maarleveld 1953; ter Wee 1962). According to the traditional interpretation, the first (oldest) phase (a) marks the maximum Saalian ice coverage, when Utrecht Ridge was built (Ruegg 1991). Each younger phase represents a readvance of uncertain magnitude during Saalian deglaciation. During each of these phases, the ice margin was highly irregular, with ice lobes extending beyond the main inland ice sheet.
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from drilling records. Where megablocks make up a substantial volume of drift-plain deposits, as in many parts of southern Alberta, the surface landform may be called a megabloek plain. Stalker (1973, 1976) was apparently the first geologist to
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106
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extend for several km, but are always quite thin (1 km wide. The megablock is nearly horizontal, and in many parts it displays remarkably little internal disturbance. Its source and distance of movement are uncertain. Bend - One of the largest megablocks yet discovered in Alberta outcrops in the bluff along the eastern side of the Oldman River about 15 km northeast of Taber. It consists of interbedded Cretaceous shale, coal, and sandstone strata. Intermittent exposures extend for >3 km along the bluff, and seismic shot-hole logs indicate the megablock extends a considerable distance behind the bluff. Maximum thickness is about 25 m with average thickness being about 10 m. Stalker (1976) estimated the Driftwood Bend megablock covers at least 10 km 2. Driftwood
The continuity and intact nature of this megablock have been stressed, but the megablock may actually include several discrete blocks lying adjacent to each other at about the same level. Some blocks are nearly fiat-lying, and others are strongly deformed (fig. 7-9). Deformation is minimal toward the southern end of the section and increases northward. At least one block near the northern end forms a large recumbent fold that is stretched southward into m61ange. The megablock is underlain by some 25 m of drift including till, and is overlain by 15 m of similar material. Immediately beneath the megablock, shear planes and slickensides are developed in the till. Once again, the source and distance of travel of this megablock are not known. Wolf Island - A megablock composed of Cretaceous shale,
Figure 7-9. Megablocks of Cretaceous strata (light tone) within thick glacial strata exposed along the bluff of the Oldman River, near Taber, south-central Alberta. A coherent megablock. The upland surface is a nearly fiat glacial plain. Note people standing at the base of the megablock, which is a shear zone in the underlying glacial sediment. B - recumbent fold with core of disrupted bedrock material resting on glacial strata. Photos by J.S. Aber (1984).
coal and bentonitic sandstone is exposed for a distance of nearly 1.5 km along the northern bluff of the Oldman River about 20 km east-northeast of Taber. It is both underlain and overlain by tills and other drift in the 75-m-high section. The megablock is more-or-less horizontal, up to 13 m thick, and is locally deformed particularly near its base, where it intermingles with underlying till. The megablock also includes up to 3 m of preglacial gravel resting on the Cretaceous strata. The gravel has a Bow Valley lithology, which is not normally found in the Oldman River system. The southern limit of Bow Valley gravel is more than 50 km to the north (fig. 7-8). Catchem - This megablock is known from only two drill holes about 20 km east of Pakowki Lake. The megablock is probably no more than 1 km 2 in extent, and its source and distance of transportation are not known.
1107
Megablocks and rafts Megablocks are seemingly ubiquitous in the Alberta Plain. The source and direction of movement for many are unknown, but some did travel considerable distance. The megablocks are not related to any single glaciation or particular topographic setting. Their emplacement appears to be widely scattered in time and space. Individual megablocks were probably transported by freezing onto the bottom of an overriding ice sheet. Thin or discontinuous permafrost would have facilitated the detachment of large, thin megablocks and their incorporation onto the base of the ice sheet. Basal melting would have later allowed separation of megablocks from the ice. At that point, megablocks could have been partly disrupted, pulled apart, folded or sheared together with drift.
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Figure 8-6. Northeastern Kansas showing Independence glacial limits, local ice movement directions, buried valleys (dotted), and locations of described sites: A = Atchison, W = Wathena. Adapted from Aber (1991, fig. 3). Reproduced from Boreas by permission of Taylor & Francis AS. Formation, Kansas; Herdla Moraine, Norway; and Systofte, Denmark. In these cases, the intrusive sediment is clay or silt rich and was injected as a result of glacier overriding. None of these structures has any morphologic expression. The Kronowo esker from Poland is a morphologic feature that was deformed in an ice-marginal setting.
Where thick Independence Formation fills preglacial valleys, as at Atchison and Wathena (fig. 8-6), its original stratification and structures can be seen in deep exposures. The Independence Formation stratotype at Atchison includes three informal members (fig. 8-7):
Atchison, Kansas The Independence glaciation was the most extensive ice coverage to take place on the Great Plains of central North America and is recognized as an important early glaciation of the Pleistocene. This glaciation is represented by the Independence Formation, a lithostratigraphic unit defined in northeastern Kansas (Aber 1991). The Independence Formation is quite old, roughly 600,000 to 700,000 years BP (Aber 1991), and is probably the oldest regionally preserved Pleistocene glacial sequence on land. Because of its great age, original glacial morphology is largely gone, and the upper portion of the Independence Formation is greatly altered by weathering and erosion. The Independence glaciation created many, scattered glaciotectonic deformations both in drift and in consolidated Paleozoic limestone and shale bedrock of the region (Dellwig and Baldwin 1965).
Figure 8-7. Photograph of Independence Formation stratotype at Atchison, Kansas. Upper and lower tills are separated by glaciolacustrine sand. I. Abdelsahed (above) and B. Nutter collect till samples from the diapir at the center of the section. Photo by J.S. Aber (1987).
Intrusions, diapirs and wedges
115
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3 km long and approximately 200 m wide. Morawski (2003) first described this esker and interpreted it as deposits that accumulated in an ice crevasse. It runs from the NE towards the SW, with a slight ENE deviation in its eastern part. This landform is represented by a row of elongated hills from several m to a dozen m high. It runs across a hummocky glacial plateau
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ii.................... i5 ............ :'::20 m thick. At culminations of individual hills, reaching 150-160 m elevation along the linear ridge of the Kronowo esker, sand-and-gravel glaciofluvial deposits are exposed on the surface, and flow till is present elsewhere on the ridge. The top of the esker, an erosional surface, is disconformably covered by a thin pavement layer of poorly sorted, massive deposits followed by a light brown sandy diamicton containing clearly visible structures indicating transport in mud flows. The thickness of this till ranges up to about 8 m in the northeastern part of the exposure. However, the thickness of the esker deposits varies from a dozen to >20 m. The deposits are represented largely by cross-bedded, poorly sorted sands and gravels. At the Kronowo exposure, the sandand-gravel esker succession is underlain by a compact till, 1.5 to 3 m thick. The till occurs only within the ice crevasse composing a basal layer of its fill. This layer lies at the same depth, or slightly lower, as the base of till composing the
Aber and Ber
122
glacial plateau surrounding the esker. The esker succession fills a tunnel cut into the Vistulian (Weichselian) till (Morawski 2003).
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Detailed analysis of the topographic trend of the Kronowo esker indicates that this is not a continuous ridge, and its individual parts are shifted relative to each other and arranged en echelon (fig. 8-21). Planes along which these shifts occur represent NW-SE oriented strike-slip faults. One of these faults shifts the quarried gravel-and-sand deposit, causing the NE quarrying works to be relocated 80 m northwards. These are glaciotectonic faults caused by NW-SE oriented horizontal stress exerted by the active ice body.
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Figure 8-21. Sketch map of the Kronowo esker: 1 - sand and gravel mantled with a thin flow till (esker sediments), 2 - till (Vistulian glaciation morainic plateau), 3 morphological axis, 4 - strike-slip fault zones, 5 - fan fold axis, 6 - pit boundaries, 7 - cross-section, 8 - borehole. Adapted from Morawski (2003, fig. 7).
Exposures in the gravel pit reveal strong compressional glaciotectonic deformations forming two fold limbs which are in contact and thrust over each other (figs. 8-22 and 23). The strike of the thrust surface is about 80 ° . The surface is inclined northwards at an angle of 60-70 °, indicating that the horizontal stress was directed from the north. This direction is also indicated by the fold asymmetry and by small, more-or-less horizontal faults observed in both fold limbs, particularly in the northern limb. The till under the esker series in lower parts of both fold limbs, north and south of its axis, lies horizontally in the original position. Towards the axis, the till layer is bent arcuately to reach a vertical position, and is inclined towards both sides. The arc diameter is about 20 m. The till, in both the exposures and boreholes, is dislocated by bending and squeezing upwards. In the southern fold limb, this break is about 10 m long; in the northern fold limb the till is truncated by a thrust surface.
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Figure 8-22. Geological cross-section of the pit walls with interpretation of glaciotectonic fold: 1 - f l o w till, 2 - till of the basal layer, 3 - esker sand and gravel, 4 - glaciofluvial sand and gravel of the esker basement, 5 - boreholes. Adapted from Morawski (2003, fig. 8).
123
Intrusions, diapirs and wedges
2
Figure 8-23. Photograph of the southern side of the fan fold with intrusive fold core to right side. Compare with left side of previous figure. Photo by Morawski (2003, fig. 11).
......
All the data collected by Morawski (2003) indicate that the original width of the crevasse in the Vistulian ice sheet was approximately 200 m. This crevasse was initially filled with diamicton, probably originating from mud flows and/or from the melted and collapsed tunnel roof. This sediment was probably subjected to short-lived erosion and subsequently deposited as till, up to several meters thick. This phase was followed by deposition of a sand-gravel succession, over 20 m in thickness, transported from NE to SW. After sedimentation and filling of the ice crevasse, probably already during the initial deglaciation, the ice sheet readvanced southwards for a short time. Only part of the ice sheet, located on the north of the crevasse, was presumably active at that time. The southern area was covered by stagnant ice. This readvance resulted in deformation and upward squeezing of the crevasse-fill (fig. 8-24, phases 1, 2). Squeezing of both the crevasse-fill and underlying deposits was initiated and took place with a contribution of vertical (loading) stress of the ice body, as was the case during formation of diapiric structures in ice-free areas (Brodzikowski 1980) and in squeezed or till-cored eskers (Rotnicki 1960). In the case of Kronowo, however, the vergence of glaciotectonic structures and the shift of individual parts of the esker by strike-slip faults, indicate a strong lateral push by the ice from the north. Such a push resulted in the formation of a large fan fold. The fold limbs were subsequently stretched and broken by expansion of the diapiric core. The limbs tilted out toward both sides and then were partly truncated by a thrust (fig. 8-24, phase 3 and 4). In the area exposed by quarrying, the distance over which the ice sheet advanced (magnitude of glaciotectonic narrowing of the ice crevasse) was determined as a sum of deformed and broken strata, and amounts to about 80 m.
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Figure 8-24. Development stages of the Kronowo crevasse infilling and glaciotectonic folding: 1 -Kronowo crevasse and reactivated ice sheet from the north (arrow indicates the direction of ice movement) initiated folding of crevasse fill; 2, 3 and 4 - successive phases of ice-sheet movement and crevasse-fill folding, crevasse width reduced to about 80 m, 5 - deposition of flow till after periodic erosion. Adapted from Morawski (2003, fig. 15).
Subsequent movement of the ice sheet resulted in en block and en echelon shifting of the crevasse-fill. The general direction of the shift was towards southeast. Individual horizontal shifts within the esker ridge can be estimated at 20-80 m. It can be approximated that local advance of the
124 ice sheet in the Kronowo region totaled a distance of ~150 m. The upper portion of the folded crevasse-fill was probably eroded by melt water. These deposits were subsequently covered by diamicton flowing down into the crevasse from melting dead ice blocks (fig. 8-24, phase 5). These observations indicate the Kronowo esker was formed by a crevasse-filling process in the continuously active ice sheet, or maybe during the initial phase of its stagnation and minor readvance (Morawski 2003). Meltwater erosion was followed by deposition, and the crevasse was infilled with flow tills and glaciofluvial sediments. The glaciotectonic fan fold was formed as a result of subsequent compression of the ice crevasse. The process of upward-squeezing of the crevassefill and partly also of the underlying glaciofluvial deposits (in a fold core) was probably caused by the same ice sheet movement which subsequently triggered strike-slip faulting across the esker ridge. The ability to define local directions of glacier movement from the spatial orientation of compressional glaciotectonic deformations is straightforward in most situations (chap. 2). However, the Kronowo esker was transected by strike-slip faults
Aber and Ber
oriented obliquely to the main folding stress. The ice mass presumably contained pre-existing crevasses (joint net), along which local movements of the ice body could occur. The ice sheet push was exerted on deposits filling a crevasse that trended obliquely to the direction of this push. Individual glacier blocks advanced not in the same direction as the main ice mass; they moved along crack surfaces, which resulted in strike-slip faults. On the basis of deformation of the Kronowo esker and comparison to regional glaciation, directional shifting in ice movement can be reconstructed. Ice advance took place from the north during the entire Main Stadial of the Vistulian (Weichselian) glaciation. Following ice retreat and stagnation, a minor readvance came from the NNW, which led to folding and compression of the esker ridge. Finally the esker was displaced along strike-slip faults trending NW-SE as a consequence of local movement of ice blocks along a crevasse network. The Kronowo exposures provide a unique opportunity for precise determination of the direction of icesheet movement during a short-lived period of its activity that occurred prior to the final deglaciation.
Chapter 9 Basement and Deep Crustal Structures Introduction An important outcome of the glaciotectonic mapping project in North America was recognition of basement and deep crustal structures as a significant category of glaciotectonism (Aber et al. 1995). Such features include basement faults and seismic zones wherein the crust responded differentially as discrete blocks to stress induced by glacial loading and unloading of the lithosphere. Such deformations are usually manifested at the time of deglaciation or shortly thereafter. Residual crustal strain may continue with decreasing amplitude and frequency long after the disappearance of ice loading. For example, an arcuate band of seismic activity is present along the Bell and Boothia arches of northernmost Hudson Bay and Foxe Channel. Basham, Forsyth and Wetmiller (1977) speculated that the Foxe-Baffin crustal block is responding to postglacial uplift independently from the rest of the Canadian Shield. The seismic band follows the edge of the Foxe sector of the Laurentide Ice Sheet. Seismic activity and block tilting toward the northeast were interpreted as the results of reactivation of preglacial structures caused by high differential stress during glacial unloading. Adams (1989) similarly suggested that much of the postglacial faulting in southeastern Canada and the adjacent United States resulted from stress release and flexural deformation of the upper crust as a result of glacial loading and unloading. The regional-scale location of faults coincides with the southern margin of the Laurentide Ice Sheet. Several zones of seismic activity are found at or near the southern limits of glaciation in stable continental crust of the United States and southeastern Canada. • Nemaha zone: northeastern Kansas and southeastern Nebraska; maximum Independence (pre-Illinoian) glaciation limit (Aber 1991). • New Madrid zone: southeastern Missouri, southern Illinois, and adjacent states; maximum Illinoian glaciation limit. • Anna zone: western Ohio; maximum late Wisconsin glaciation limit. • Atlantic coastal zone: from northern New Jersey to southern New Brunswick; maximum late Wisconsin glaciation limit. • Saint Lawrence zone: northern New York to southeastern Quebec; deglaciation ice limit of 12,000 years BP (Dyke and Prest 1987b).
These seismic zones lie at the intersections of northeasttrending (late Proterozoic) and n o r t h w e s t - t r e n d i n g (Mesozoic) basement structures, and the seismicity of these zones is related to continued opening of the North Atlantic Ocean (Barosh 1990). Aber et al. (1993) speculated that earthquakes in these and other ice-marginal seismic zones were accentuated by glacial loading and unloading. Other crustal disturbances one km or deeper in the subsurface also are included in this category. Incompetent strata, such as salt, were mobilized under the impact of glacial loading at considerable depth. As an example, White (1992) documented displacement of salt south of Lake Erie in northeastern Ohio (fig. 9-1). The combined thickness of Paleozoic salt beds exceeds 60 m in the region south of Cleveland near the limit of glaciation. The thickest salt (>90 m) is found in narrow zones within the cores of doubly plunging folds. The fold axes trend generally parallel to the south shore of Lake Erie and to the glacial limit. Near the limit of glaciation, the folds become somewhat arcuate and are spaced closer together. White concluded that stress imposed by successive glaciations in the Erie basin had mobilized the salt southward. The following examples include basement faults and seismic zones from the Canadian Shield and Appalachians in North America and the F e n n o s c a n d i a n Shield of northern Scandinavia in Europe. Deep salt displacement is examined in the Finger Lakes district of New York, and buried faults are described from Poland.
Canadian Shield and northern Appalachians This class of glaciotectonic structures is widely documented in the northern Appalachian region and from the southern Canadian Shield. The structures are of two types: small faults in basement rocks and zones of seismic activity. The small faults included are those considered to be the direct results of glacial loading and unloading, in contrast to faults resulting from regional tectonics, mine pop-ups, or frost heave. Glaciotectonic basement faults are found primarily in Paleozoic slates of the Appalachian region (Chalmers 1897; Goldthwait 1924) and in Proterozoic slates of the Canadian Shield of western Ontario (Lawson 1911; Oliver, Johnson and Dorman 1970). Small faults are also reported from other regions and from well-consolidated argillite, sandstone and limestone (Cushing et al. 1910; Broster and Burke 1990).
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multiple Pleistocene glaciations, the region north of the Valley Heads moraine was covered many times, as evidenced by the magnitude of glacial landscape modification. However, only rarely during maximal phases did glaciation extend into the Appalachian Plateau beyond the Valley Heads moraine. The Ontario Lowland and northern margin of the Appalachian Plateau were, therefore, covered by relatively thick ice more frequently and for longer periods than was the region south of the Valley Heads moraine (Cadwell and Muller 2004). The cumulative effect of repeated differential loading by glaciation was migration of salt away from the northern Finger Lakes vicinity and toward the zone of lesser ice loading south of the Valley Heads moraine. Depth of salt displacement reached 1-2 km in south-central New York and northeastern Pennsylvania. A related structure is a salt-cored anticline that lies beneath and parallel to the N-S axis of Lake Seneca (fig. 9-13). Glacial erosion removed considerable overburden from this longest and deepest of the Finger Lake valleys. Upon deglaciation, salt flowed toward this zone of lower overburden pressure (L.E Dellwig, pers. com.).
The anomalous distribution of salt is restricted to the Finger Lakes region and does not extend eastward or westward. Farther south, in northeastern Pennsylvania, thickened salt cores are found in overturned and thrust folds of the Appalachian Plateau >2 km deep. The direction of overturning and thrusting is to the southeast, toward the Appalachian Mountains. White (1992) interpreted these anomalies as results of massive salt displacement southward in the downdip, downglacier direction under the impact of glacial loading over the Ontario Lowland region.
Poland contains several excellent examples of deeply buried basement structures that were reactivated during repeated glacial loading and unloading, and may have contributed to surficial glaciotectonic deformations as well. These structures have been documented through extensive test drilling and geophysical investigations.
Ice thickness in the Ontario Lowland region was on the order of 1500-2000 m, a mass load considered sufficient to induce halokinesis (White 1992). In contrast, ice thickness over the Appalachian Plateau south of the Valley Heads moraine was likely no more than a few 100 m. During the course of
Buried faults, Suwalki Lakeland m The Suwatki Lakeland is situated in northeastern Poland bordering with Russia (Kaliningrad district) and Lithuania (fig. 9-14). Analysis of geological data obtained from deep boreholes drilled in this area during past half century shows clear influence of the
Polish basement structures
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Figure 9-13. Wide-angle view of Seneca Lake looking toward the north from near the southern end. It is about 2 km wide at this point and extends ~50 km northward. A salt anticline trends under the lake trough axis. Photo by J.S. Aber (2005).
134
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Figure 9-14. Tectonic map of the Suwatki Lakeland district in northeastern Poland with glacial surficial features. 1 - tectonic discontinuities, 2 - main faults, 3 -Suwatki Anorthosite Massif (SAM), 4 - Wi~.ajny Elevation, 5- glacial features of phases and subphases of the Vistulian Glaciation, 6 - glacial tunnel valleys, 7- zones of glaciotectonic deformations. After Ber (2000, fig. 3; modified).
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Figure 9-15. Schematic model for possible structural displacement and glaciotectonic thrusting associated with glacial loading and unloading of a deep-seated fracture in the substratum. In stage B, fault movement induced by glacial loading has created a buried obstacle over which ice-shoved hills are thrust. Adapted from (Liszkowski 1993, fig. 7). geological structure and tectonic style of the crystalline basement on the sedimentary cover, including unconsolidated Quaternary deposits and glaciotectonic features (Ber and Ryka 1998; Ber 2000).
gravity and seismic investigations, as well as maps of the sedimentary cover. Depth to the crystalline basement ranges from only 0.5 km in the northeast to more than 1.6 km in the northwest of the Suwatki Lakeland.
Geological data concerning the crystalline basement (Kubicki and Ryka 1982; Graniczny 1998) are based on results of
Since late Proterozoic time, the crystalline basement of this region has been subjected to continuing and repeated stress
135
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Figure 9-16 (left). Simplified section across the SzczytnoOkrqgte Lake dislocation zone (SJO-DZ) in northeastern Poland and adjacent Lithuania. 1 - boreholes terminated in: a - crystalline basement, b - sediments of Lower Paleozoic; 2 - s e d i m e n t a r y cover; 3 - crystalline basement; 4 sedimentary continuities (a) and angular sedimentary discontinuities (b); 5 -fractures and faults in crystalline b a s e m e n t (a) and o f Caledonian age (b); 6 - ;:one o f glaciotectonic deformations in Quaternary sediments; 7 movement direction of Pleistocene ice sheets; L-P ice sheet limit of the Leszno-Pommeranian Stadial of the Vistulian Glaciation; W, Pm - subphases (Wigry and Pommeranian) ice sheet limit; AR- P R - Archaic-Precambrian; • - Cambrian; 0 - Ordovician; S - Silurian; D - Devonian; P - Permian; T - Triassic; J - Jurassic; Cr- Cretaceous; Pg-Ng - PaleogeneNeogene; Q - Quaternary; S J O - D Z - zone o f tectonic discontinuity Szczytno-Okrqgte Lake - Deep-seated Zone. Adapted from Ber (2000, fig. 8).
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occurred several times in the Suwatki Lakeland. Such crustal movements were often focused along pre-existing fault zones and left an imprint in surficial glacial geomorphology (fig. 9-15). It can be assumed that loading and relaxation within the crystalline basement occurred several times, during each glacial and interglacial period, causing rhythmic glaciostatic movements. The amount of vertical depression during icesheet advance was up to 200 m, depending upon the ice thickness and duration. In interglacial periods, relaxation and complete rebound of the crust took place. The ice sheets from earlier, pre-Vistulian glaciations were thicker and more widespread. During their advances they crossed latitudinal fracture zones where ice-laid deposits were subjected to glaciotectonic deformations (figs. 9-14). It cannot be precluded that the process of ice sheet retreat could depend partly on the structural pattern of the deep basement, i.e.
conditions that led to formation of new and reactivation of pre-existing meridional (N-S) and latitudinal (E-W) tectonic fault zones (Znosko 1984). North-trending faults played a special role within the Suwalki anorthosite massif (SAM, fig. 9-14). They divided the area into low and high tectonic blocks already in the late Proterozoic. Tectonic movements of the crystalline basement were superimposed by glaciostatic loading and unloading, which took place during glacier advance and retreat (Weertman 1961; M6rner 1977, 1980; Liszkowski 1993). The former resulted in crustal depression and the latter in rebound, which P~TIC m
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Figure 9-18. Sketch map of glacial landscape features in the Suwatki Lakeland district of northeastern Poland showing festoon and valley-side glaciotectonics. Adapted from Ber (2000, fig. 36).
variable activity along meridional and latitudinal faults separating individual tectonic blocks and differential shifting of the blocks. Glaciostasy, operating along dislocation zones and faults, exerted in turn an essential influence on the Vistulian glaciation in the Suwa~ki Lakeland and throughout northeastern Poland. Vertical movements along latitudinal and near-latitudinal faults conditioned development of parallel-trending river valleys, topographic lows and highs, maximum limits of ice sheets, and recession and oscillatory ice-marginal zones, where glaciotectonically deformed moraines are observed (Liszkowski 1993; Ber 2000). Zones of glaciotectonic deformations are associated with basement faults and vast stepwise tectonic fault zones; for example, the Szczytno-Okr~gte Lake dislocation zone (SJO-DZ, fig. 9-16) varies from a width of 10s of meters (northeastern Poland) up to >100 km in Lithuania (Sliaupa 1996). In general, north-, NNE- and NNW-trending (meridional) faults exerted an influence on the formation of similarity in orientation of subglacial and surface crevasses in the ice body, resulting in the development of eskers and subglacial tunnel valleys. A similar situation existed in Estonia (Tavast 1998); conversely in Belarus, eskers were controlled by latitudinal faults (Karabanov 1997). Pleistocene deposits of the Suwatki Lakeland were subjected to glaciotectonic deformations due to advance and retreat of
Figure 9-19. Main configuration of Neogene and Quaternary substrates and mid-Pleistocene horizons in western Mazury and Warmia, northeastern Poland. 1 - depressions of Neogene substrate, 2 - depressions of Quaternary substrate, 3 - area occupied by Holsteinian Sea, 4 - area occupied by Eemian Sea, 5 - main streams of Mazovian Interglacial, 6 main streams of Eemian Interglacial. Adapted from Marks (1988, fig. 19).
successive ice sheets. Deep-seated faults manifested themselves as flexures in glacial and glaciofluvial deposits. Through the whole Pleistocene, glaciostatic movements operated along meridional (N-S) and latitudinal (E-W) dislocation zones and faults. The most legible record of this effect is visible in the present-day relief of the Suwatki Lakeland, which was conditioned by ice streams, lobes and tongues that followed in turn patterns of the crystalline basement and its sedimentary cover, including Pleistocene deposits (Ber 1987, 2000; Ber and Ryka 1998). The consistency of the main orientations and present-day relief
137
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with the tectonic pattern of crystalline basement (fig. 9-14) demonstrates that pre-existing tectonic zones were rejuvenated in the Pleistocene due to differential crustal movements, which is also confirmed by photolineament analysis (Graniczny 1998). The Wi2ajny Elevation, located in the northern part of the Suwatki Lakeland, had a special influence on variability and movement directions of the Vistulian ice sheet, and also older ice sheets. The ridge developed on the pre-existing paleostructure composed of Cambrian and Valdai deposits. Owing to the tectonic situation through the whole Pleistocene history, it was an isolated basement high, similar to those observed in other areas of northeastern Poland (Szeskie Hills, Elbl~g Upland, Gdrowo Hills), in Belarus (Grodno Upland), and Lithuania (~emajtija Upland), as well as Latvia and Estonia. All of these are characterized by strong glaciotectonic deformations of surficial strata (fig. 9-17). The Vistulian ice sheet caused deformation of the basement, inducing static (vertical) stress and dynamic (horizontal) stress, as presented in a model by Rotnicki (1976). The stresses resulted in reactivation of tectonic fracture zones and mainly latitudinal faults controlling ice sheet extents and
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stagnation periods during recession (fig. 9-14). This gave rise to the formation of various ice-marginal landforms, of which festoon relief and valley-side glaciotectonics are distinctive examples (Ber 1987, 2000; fig. 9-18). Stressinduced deformation activity of the ice sheet, as well as the formation of festoons, occurred mainly in areas composed of till. It may be possible that development of glaciodepressions and glacioelevations, forming large festoons and lobes, was associated with mobility of individual tectonic blocks of the crystalline basement. Warmia and western Mazury Lakeland w The present relief of the Warmia and western Mazury Lakeland (northeastern Poland) was shaped by glaciotectonics and neotectonics connected with the structural plan of the old basement and activated along meridional (N-S) and especially by parallel (E-W) fractured zones. These zones were formed at the lithologic boundaries between Precambrian structural units of different density. Especially, the parallel geological boundaries of the crystalline basement and large faults connected with them had a basic effect on the parallel direction of many glacial forms of the recent landscape, such as marginal zones of the Vistulian Glaciation, hills of end
138
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Figure 9-22. Isolated Pleistocene uplands (cupola hills) in comparison to pre-Arenigian (Ordovician) paleo-structures in northeastern Poland--Mazury, Warmia and Suwatki districts, showing their influence on direction of ice-sheet advance and glaciotectonic disturbances during the Vistulian Glaciation. 1 - positive pre-Arenigian structures, 2 - isolated Pleistocene uplands (cupola hills) glaciotectonically disturbed, 3- interlobate zones which influenced directions of ice movement, probably tectonically predisposed to glaciotectonic disturbance, 4 - direction of ice advance, 5 - maximum limit of Vistulian Glaciation. Adapted from Ber (2000, fig. 39).
and push moraines, and subglacial channels. Parallel boundaries likewise influenced zones of glaciotectonic deformations (Marks 1988; Ber 2000; Morawski 2004a). In the Warmia region, according to Marks (1988), the surface of the Quaternary substrate has a system of elongated elevations and depressions with general NNW-SSE direction, which are associated with features of the Neogene substrate and depressions and faults in the top surface of Cretaceous rocks (fig. 9-19). Similar direction (NNW-SSE) is marked in the system of river valleys of the Mazovian (Holsteinian) Interglacial (see fig. 9-22). However, the fault zones within the Mesozoic complex, known from vicinity of Lidzbark Welski (fig. 9-20) and ~uromin (fig. 9-21), showed at the same time a gradual vanishing of faults downwards and a change of throw or of its direction. This demonstrates that fault zones have been rejuvenated during the Pleistocene many times (Marks 1988). The isolated uplands in the present landscape in this area (Elbl~g Upland, G6rowo Hills and Dylewska Hill) contain strongly disturbed glaciotectonic internal structure and correspond with buried positive pre-Arenigian paleostructures composed of Cambrian and Valdai deposits,
named by Kotafiski (1977)" Elbl~g Elevation, G6rowo Itaweckie Hump and Dylewska Bank (fig. 9-22). In the present landscape of the southwestern Mazury, the Lubawa Upland with its highest peak of Dylewska Hill (312 m elevation) played a significant role. Dylewska Hill is composed of upthrusted Paleogene, Neogene and Quaternary sediments (Marks 1978; Gat~zka and Marks 2000) and corresponds with a pre-Arenigian high in tectonically predisposed deep basement. To the east of Dylewska Hill, the Quaternary substrate is disintegrated into some elevations and depressions, running NNW-SSE and highly concordant with the deeper, pre-Arenigian paleostructures. This coincidence connected with the activation of fault zones, according to Marks (1988), affected palaeogeography of the area during the Pleistocene, especially since the Mazovian (Holsteinian) Interglacial (see figs. 9-19 and 9-22). In the western Mazury Lakeland, deeply rooted glaciotectonic thrust structures, reaching a depth of about 300 m, have been found near Ortowo (Morawski 2004a). Lateral extent of these structures is up to 5 km and Oligocene through Pleistocene (Odranian Glaciation) deposits are involved. The maximum gradients of the thrust surfaces range from 15 ° to 20 ° and
139
Basement and deep crustal structures
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the thrust vergence is towards NE, opposite to the general direction of ice sheet movements in the Polish Lowlands. The thrusts, according to Morawski (2004b), seem to follow a structural slope, probably formed in the place of the buried faults occurring in the topmost Cretaceous through Miocene strata, and they developed probably as a result of vertical pressure exerted by the ice on the deeper basement. The main structural elements of the glaciotectonic deformations are glaciotectonic thrusts (slices) and diapiric upthrustings (see fig. 2-19). As Morawski (2004b) described it, slices thrust over each other, create a contractional duplex, with the vertical component of the stacking process being a derivative
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140
Aber and Ber
of pressure, which took place prior to or during Odranian (Drenthe) Glaciation.
NW ODOLAN6W BASIN --~
Ostrzesz6w Hills, Wielkopolska Lowland ~ The Ostrzesz6w Hills, situated at the southeastern end of the Wielkopolska Lowland, are the easternmost and highest part of the Silesian Rampart, which includes several other major glaciotectonic uplifts, such as Trzebnica Hills (fig. 9-23). Altitude of Ostrzesz6w Hills exceeds 280 m. Ostrzesz6w Hills extend 34 km N-S, are 4 to 10 km wide, and display an arcuate curvature, visible on Landsat images (Aber, Ruszczyfiska-Szenajch and Krzyszkowski 1995), concave toward the west (fig. 9-24). The area extending toward the west and northwest from Ostrzesz6w Hills is the Odolan6w basin with its bottom at the level ~120 m above sea level, which marks surficially the region of the buried depression within the Neogene substratum.
Rotnicki (1967) interpreted Ostrzesz6w Hills as a thrust end moraine, pushed from the west by the Wartanian (Saalian) Glaciation. Internal structure consists of steeply dipping imbricated blocks, pushed from the west, that include Pleistocene, Pliocene and Miocene sediments. West of Ostrzesz6w Hills the glaciodepression of Odolandw basin marks the source region for material thrust into the hills (fig. 9-24). Odolan6w basin and Ostrzesz6w Hills are localized within the dislocation zone of block structures above a horst within hard Triassic bedrock, beneath unconsolidated Neogene sediments (fig. 9-25). According to Markiewicz and Winnicki (1997) this situation is most important for the reconstruction of the factors stimulating the origin of deep glaciotectonic deformations occuring in this area, which were favored by diversified geological structure of the Neogene upland with positive (horst) structures of hard substratum (fig. 9-26). Glaciotectonic movements of the Triassic substratum probably were connected with processes resulting from irregular ice loads exerted repeatedly during multiple glaciations. Dislocations of block structures were influenced additionally by halotectonic movements of the underlying Zechstein salt beds, again enhanced by multiple glaciations.
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Chapter 10 Applied Glaciotectonics Introduction
Glaciotectonic deformation produces locally severe disruption of the normal geology and terrain, and therefore its recognition and structure are important for many human activities. The principal deleterious results of glaciotectonism are disruption and deformation of the stratigraphic sequence, reduction of sediment or rock strength, and major increase in substratum variability. Consideration of glacially thrust terrain is important for mine planning and operation, drift prospecting and mineral exploration, all manner of construction, soils mapping and utilization, and all aspects of ground-water development and protection (JCrgensen 2005). Glaciotectonic fracturing of bedrock presents special problems for mining and civil engineering compared to similar unfractured substratum. At the Heath Steele Mine in northern New Brunswick, for example, glaciotectonic fractures were implicated in failure of part of the mine wall. The mine was opened in a Pb-Zn ore body within porphyry and schist bedrock in the autumn of 1990. In the autumn of 1991, part of the 25-m-deep open-pit wall suffered a series of landslips and began to collapse, and the mine was closed (backfilled) the following year. The wall failure was attributed to a combination of circumstances (Park and Broster 1996). • Glaciotectonic fractures and reactivated joints with clay and sand fillings. • Presence of highly permeable and weathered joints that allowed water penetration. • Increased fluid pressure following several days of heavy rainfall. Glaciotectonic terrain may be expected to be texturally, lithologically, geochemically and geotechnically anomalous compared to the surrounding terrain. The terrain may provide data on the local subsurface stratigraphy, because thrust masses may include sediment and bedrock shoved up from the substratum. The depression upglacier from a thrust mass, if present, may serve as a window into the underlying stratigraphy. Ice-thrust terrain often has thinner drift cover than surrounding terrain, and in some places the bedrock itself may be exposed. Folding may produce inversion of stratigraphy and faulting may lead to repetition or omission of strata, that if unrecognized would cause problems during interpretation of drilling data.
Problems in soil mapping over glaciotectonic terrain originate from the high lateral variability of the deformed substratum. The parent material, which is of major importance for soil mapping, may change repeatedly from sandstone to shale, chalk, till, clay, or gravel over a few 10s of m. If bedrock or other subsurface sediments contain harmful substances, thrusting can bring these materials into the pedogenic zone. For example, bedrock thrust to the surface can increase soil salinity or acidity, because of its geochemistry compared to the normal till cover. One aspect of glaciotectonism often overlooked in applied studies is the depression left at the source of thrusting. This could be important during mineral-exploration and mineplanning phases of resource development. The false assumption that an ore body continues across a glacial depression would yield overly optimistic estimates of total reserves. Sudden discovery of the absence of ore during mining could cause disruption of the mining schedule with potentially serious economic implications. Use of a combination of test drilling and shallow subsurface geophysics between test holes is the most economical way to avoid a surprise discovery of this kind. In some coal mines of western Canada, glaciotectonism has locally removed the coal from areas up to 1 km 2. At one mine, ice thrusting extended deep enough to remove about 12 million m 3 of coal leaving a depression that is partially masked by the cover of later Quaternary sediment (fig. 101). With test holes drilled only at positions A, B, C and D, the interpretation would be that the coal subcrop is continuous. Additional drilling, such as holes E, F and G, would indicate that coal is missing. The transported coal remaining downglacier from the depression is of no use for mining, because the coal exists in small hills close to the surface. This position means the coal has been oxidized, so that its calorific value is reduced greatly and the moisture content increased. The following examples are situations where glaciotectonic disruption of normal bedrock stratigraphy and structure has created serious hazards or economic consequences for human activities. The examples are from large- and small-scale, open-pit mining operations, respectively in Alberta, Canada and Fur, Denmark, and from highway construction through ice-thrust ridges in Saskatchewan, Canada. They illustrate the importance of recognizing the special substratum conditions in glaciotectonic terrain.
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Highwall failure, Highvale coal mine, Alberta Highwall failures in glacially thrust bedrock have occurred in several coal mines of western Canada, including the Highvale Mine west of Edmonton, Alberta (fig. 10-2). The mine was owned by TransAlta Utilities Corporation, and production exceeded 11 million tons per year. The adjacent Sundance Power Plant was the largest electrical generating station in western Canada (Tapics 1984). Open-pit mining here involves two basic operations" 1) removal of the overburden above the coal using a dragline
Figure 10-3. Schematic cross section of typical open pit mine showing overburden bench and highwall. Taken from Aber, Croot and Fenton (1989, fig. 8-4). (fig. 10-3), and 2) removal of the coal seam using power shovels (fig. 10-4). The dragline operates from a bench above the coal, removing the overburdern above and below the bench and casting it onto the spoil pile beyond the current mining pit. Structural integrity of the bench and highwall is a crucial factor for successful mining. Glacially thrust bedrock is weaker than undeformed bedrock, and as a result highwalls and benches cut into this material have a greater tendency to fail. The implications of highwall failure are serious; temporary benches excavated in the highwall serve as transportation corridors and as working surfaces for heavy equipment. In addition to the risks posed to men and machines, such failures also result in potentially sizable extra costs for rehandling overburden material, disruption of mining schedules, and outfight loss of minable coal.
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Figure 10-4. Power shovels removing the coal seam at the Highvale Coal Mine, Alberta, Canada. Photo by J.S. Aber (1984). The regional bedrock geology consists of flat-lying, upper Cretaceous and Paleocene, non-marine, coal-bearing strata. The coal belongs to the Ardley Coal Zone of the Scollard Member of the Paskapoo Formation (Carrigy 1970; Irish 1970; Holter, Yurko and Chu 1975). Six distinct and laterally continuous seams are mined. They are separated by shale and bentonite partings, with a cumulative coal thickness of about 10 m. Quaternary strata consist of a discontinuous cover of glaciofluvial and glaciolacustrine sediment over till (Andriashek et al. 1979). Glaciotectonism of the bedrock is extensive, although most disturbed sites are 12 m.
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The classic study of glaciotectonic disturbance of mo-clay was published by Gry (1940, 1979), who applied methods of structural geology for analyzing the ice-push deformations. Major structures are large, rootless folds of mo-clay with or without deformed glacial strata. Disturbed mo-clay is present in composite ridges that make up the northern third of Fur (fig. 10-17). The ridges reach a maximum height of 75 m and define a gentle arc, concave northward. Close agreement exists between structural and topographic trends along the arc, which suggests the composite ridges were deformed by a northerly ice-lobe advance coming across the Limfjord basin.
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basis of gross lithology and industrial usage into four informal series. • Upper series - almost ash-free diatomite with thin ash layers (+ 119 to + 140), minor industrial use, about 8 m. • Ash series - thin diatomite layers interbedded with many, thick ash layers (+1 to +118), no industrial use, about 17 m.
The unusual properties of mo-clay have several industrial applications, so it has been quarried for many decades at various locations, principally on northern Mors and northern Fur. Primary use is for ceramic products. Structural disturbance of mo-clay is the main concern for planning and developing quarries. Quarries generally follow the strike of folded strata until those strata terminate against a fault or the overburdern becomes too thick.
Difficulty was encountered in excavation at the ManhCj quarry (fig. 10-18), during the early 1980s, as a result of irregular pockets of sand and gravel included within the mo-clay. This sand and gravel contaminates the mo-clay and diminishes its quality as a raw material, which has significance for ultimate extraction of the estimated mo-clay reserve and for day-to-day mining economics. The Geological Survey of Denmark was commissioned to investigate the problem and make recommendations to the quarry operators (Pedersen and Petersen 1985). The middle portion of the Fur Formation, totaling about 35 m in thickness from just above ash + 19 to the upper part of the claystone series, is exposed in the quarry. The portion between ash layers -13 to -19 is the quarry interval (see fig. 10-16). Normal thickness of the quarry interval is increased at this location as a result of reverse faulting. The overall structure consists of folds trending NW-SE with a culmination near the center of the quarry. Fold axes plunge about 20 ° SE in the eastern part of the quarry, whereas folds plunge about 15 ° NW in the northwestern part. This corresponds closely to the original land topography. Glaciofluvial sediment, consisting of sand and pebble- to cobble-sized gravel, fills channels cut into the folded moclay and in places contains reworked masses of brecciated mo-clay. The channels follow syncline troughs and are cut
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syncline troughs in the glacial foreland, and fissures opened on anticline crests. Anticlines became overturned with continued deformation, and sediment filling of the open fissures with glaciofluvial sand and gravel took place. Thrusts developed along anticline axial planes and cut up into channel fills. The crests of
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Figure 11-17. Cross section of Trzebnica Hills, southwestern Poland. Note imbricate stacking of glacial strata and preglacial (Miocene) sediments above the edge of a buried horst (uplift) in hard Triassic bedrock. Based on unpublished information from Krzyszkowski (1993). Compare with Ostr~esz6w Hills (fig. 9-25).
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~
shoved hill, and basal till may build up on the proximal side. Such features would ideally outline the frontal and lateral margins of the ice lobe at the time of thrusting and may be preserved in those cases where ice did not later overrun the hill. Continued glacier advance may eventually overrun the iceshoved hill, at which point a penetrative, highly sheared, metamorphic style of deformation may develop beneath the ice (Hart and Boulton 1991). Erosion of the hill provides
Q
2
4:
6
8
~iO
reworked sediment for a discordant till cover, and gradually a cupola-hill morphology develops. Further subglacial modification could produce streamlined, crag-and-tail, drumlin forms (Moran et al. 1980; van den Berg and Beets 1987). The ice-shoved hill ultimately may be destroyed completely in the subglacial environment. Under ideal circumstances, a series of ice-shoved hills may be created at different stages during ice-lobe advance and manage to survive throughout the glaciation. This results in
1
6
4
A
b
........
e
r
,.
and Ber
Figure 11-18. North-south cross section between Schoonebeek and Hengelo in the eastern Netherlands showing position of ice-shoved hill and basin at point where Tertiary strata rise toward the surface and Pleistocene sediments become thinner. Ice movement from north; note large vertical exaggeration. Adapted from van den Berg and Beets (1987, fig. 6).
Slcmbols: .....
'De¢o||emenl
~~; ~
High pore~wate:r
................~
-~""/
presSUre
~
Bas~l T|II
~
ice-con!L=¢t D~It
._.i/~
a glaciotectonic landscape (fig. 11-20), in which narrow belts or loops of ice-shoved hills alternate with wide, low basins. This situation is seen in the Netherlands, for example, where multiple stages of ice-shoved hills are found (see fig. 5-32). Glaciotectonic landscape is also well developed in northeastern Poland (fig. 11-21 and 22). The lobate pattern of glaciation and resulting glaciotectonic phenomena vary greatly in their types and geomorphic expressions. Nonetheless certain basic elements may be expected in subglacial, ice-marginal, and proglacial positions relative to the ice lobe. Melt-water drainage is a key factor for the reconstruction of paleo-ice lobes, as depicted by characteristic geomorphic forms: tunnel valleys, eskers, spillway channels, kames, outwash fans and deltas, etc. Such melt-water features are an integrated aspect of the lobate pattern of glaciation and are often closely connected with glaciotectonism. Glaciotectonic structures and landforms are, thus, part and parcel of the overall ice-lobe and melt-water impact on the landscape. Based on representative lobes of the southern Laurentide ice sheet, Clark (1992) recognized two principal types of ice lobes and their associated geomorphic features for the late Wisconsin glaciation. • Relatively thick lobes with steep margins, such as the Green Bay lobe (fig. 11-23). These lobes presumably advanced over permafrost by a combination of plastic flow and limited basal sliding. Such lobes were able to maintain continued movement for an extended period of time, which resulted in well-developed drumlin fields behind end moraines. Subglacial melt-water features, such as eskers and tunnel valleys, are generally lacking.
Figure 11-19. Schematic model for proglacial thrusting and subglacial modification of an ice-shoved hill during ice-lobe advance. A - initial proglacial thrusting, B - building icemarginal composite ridge, C - overriding and smoothing cupola hill. Note creation of ice-scooped basin (hole) behind the ice-shoved hill. Not to scale; subglacial melt-water features not shown. Modified from aber (1982, fig. 3).
• Relatively thin lobes with gentle margins, such as the James lobe (fig. 11-24). These lobes advanced rapidly over thawed ground by surging on water-lubricated beds or deforming beds. Repeated advances alternated with ice stagnation; hummocky moraine, eskers, tunnel valleys and other geomorphic features of stagnant ice are common.
Distribution of glaciotectonism
165
................
.~.~.~.~...,~. ...... ........
Figure 11-20. Long profile of idealized glaciotectonic landscape during ice-lobe advance and overriding. Note creation of ice-scooped basins between ice-shoved hills. Not to scale; adapted from Aber (1982, fig. 4).
° -
, %k,.._
. ~
f~
a
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~.~
'
?:k~
? .........................................
In our experience, conspicuous constructional glaciotectonic landforms are mostly associated with ice-marginal positions for the latter type of ice lobe. Large hill-hole pairs, composite ridges, and cupola hills are characteristic of relatively thin
Figure 11-21. Sketch map showing festoon pattern of iceshoved ridges (a) and intervening basins (b) for the Suwatki Lakeland, Poland. Adapted from Ber (1987, fig. 6).
Figure 11-22. Festoon pattern in the landscape in vicinity Elk Upland, northeastern Poland. Shaded relief image derived from digital elevation model. Image courtesy of S. Ostaficzuk (2005). Scale in km.
22~55
22,6
22.65
22.7
22~75
22,8
22,85
22.9
1
6
6
A
i
R~
8~ o
~
e
Figure 11-23. Major moraines and drumlins of the Green Bay ice lobe (A), reconstruction of ice-surface morphology (B), and profile of ice margin (C). Green Bay lobe is located in eastern Wisconsin, an area underlain by well-consolidated Paleozoic sedimentary bedrock. Adapted from Clark (1992, fig. 9).
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27,000: t,~:MOO to 31,400 ±: 2000
In coastal cliffs on Wolin Island, glaciotectonic deformations are connected with the following sedimentary deposits: Cretaceous marls, forming scales and rafts; grey, Middle Polish (Grodno) till and brown till (Trz~sacz) of Late Vistulian age; and intertill sands and silts with gravels (fig. 11-39). In the bottom of tills, numerous glaciotectonic structures are present (Bor6wka, Goslar and Pazdur 1998). Wolin Island displays a distinct connection of surface landforms (morainic massifs) and their longer axes of elevations and glaciodepressions to azimuths of the main glaciotectonic structures. The surface of Wolin Island is diversified by numerous ridges of push moraines built with glaciotectonically disturbed Pleistocene deposits including Cretaceous rafts. According to Bor6wka, Goslar and Pazdur (1998), the geologic and geomorphologic data suggest that the glaciotectonic structures present in the cliff section are connected with the retreat (not advance) of the last (Vistulian) ice sheet. Vistulian glacier flow direction and deformation of Wolin Island deposits traditionally were interpreted as from NNENE directions. However, Lagerlund (1987, 1997; et al. 1995) and Bj6rklund, Lagerlund and Ing61fsson (1998) challenged this model. According to them, western and central portions of Wolin were deformed by ice movements from the west and northwest, which distinctly deviate from the predominating NNE-NE directions. Similar anomalous ice movement directions in Sweden (western Sk~ne) and northeastern Germany (Usedom, Riigen) are indicated (fig. 11-40). In eastern Wolin, in contrast, glaciotectonic structures suggest ice pushing from the northeast. Belarus m Only the northern part of the country is located within the limit of Poozerie (Weichselian) glaciation (fig. 11-41). The maximum limit of this glaciation is marked mainly by push moraines with glaciotectonically disturbed
Grodno till M i l l e Polish
Figure 11-39. Generalized stratigraphic sequence of Pleistocene deposits resting on dislocated Cretaceous strata from Wolin Island. Ages given in radiocarbon years before present. Adapted from Bor6wka, Goslar and Pazdur (1998, fig. 1).
~11~-. Group 2 .,~---GK~up 3 ~ . G r i p 4: -.~ Group 5 . ~ ....... G:ro~up6 100kin
5
Bedin
6 Po~an
Figure 11-40. Anomalous directions of Weichselian ice movement in the vicinity of Riigen (R) and Wolin (W) islands. Arrows show groups of anomalous directions at different sites and different stratigraphic levels. Adapted from Lagerlund (1997, fig. 1).
177
Distribution of glaciotectonism
internal structures composed of dislocated tills and glaciofluvial deposits. According to Karabanov (2000), folded (Sopotskin) and injective ridge systems (Ushahi Lake group) are connected with faults in the crystalline and sedimentary basement of northern Belarus. Most common among glaciotectonic structures are folded dislocations, glaciodepressions, diapiric hills and ridges, and rafts composed of Cretaceous, Paleogene or Neogene strata. Glaciotectonic landforms often are associated with dislocations of Cretaceous rocks (e.g. Sopotskin, Grodno). The Grodno Highland area is a well-known example of a glacial inter!obate massif (cupola hill), similar to insular glaciostructural-accumulative highlands in Latvia and Estonia or the isolated Pleistocene elevations in Poland. It was influenced by the main ice streams during Middle and Late Pleistocene glaciations (Karabanov 1987). According to Pavlovskaya and Karabanov (2002), most of Grodno Highland is deformed in the shape of hill-hole pairs, stretching generally in W-E direction, with impressive dislocations of Cretaceous rocks at Grandichi, Pyshki, Melovaya Gora and Sopotskin. The chalk quarry at Pyshki, for example, reveals folded scales of upper Cretaceous chalk and chalky marl deposits, as well as upper Miocene and lower Oligocene (glauconite and quartz, sand and silt) strata, and injective forms and megablocks (Karabanov 2002a). These deposits are overlain by Pleistocene till and glaciofluvial sand (fig. 11-42). The Cretaceous deposits are normally located 80-100 m below water level in the Neman River, but where involved in glaciotectonic dislocations they are exposed at the surface.
i
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i
.....
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Y
[............................. ~:2000 m thick. Sea-floor drilling revealed the upper 20-25 m of bedrock is
Several representative situations serve to illustrate the nature and conditions of glaciotectonism, beginning with the
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inferred ice sheet limits submerged shorelines
Figure 11-58. Bathymetry of the Pechora Sea floor showing three glacial limits during late Weichselian glaciation. The Kurentsovo Line contains large ice-thrust bedrock ridges formed by an ice sheet derived from the Barents Sea. Takenfrom Gataullin, Mangerud and Svendsen (2001, fig. 4); reprinted with permission from Elsevier.
Distribution of glaciotectonism
189 I
r
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Figure 11-59. Geologic section of strata and structures exposed in cliffs and recorded in boreholes along the Ob River near Atlym, western Siberia. Symbols: 1 - glacial strata, 2 - lower Miocene to upper Oligocene terrestrial silt, 3 - middle Oligocene terrestrial sand, 4 - lower Oligocene to Eocene marine clay, 5 - Eocene diatomite, 6 - Paleocene marine clay, 7 - faults observed (a) and inferred (b), 8- borehole. The profile is ~15 km long; adapted from Astakhov, Kaplyanskaya and Tarnogradsky (1996, fig. 16). Copyright John Wiley and Sons Ltd. Reproduced with permission.
t 2.~
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-~.
Loc~ions o stri~,ed boulders
:~
D i s p I a c e d ( f o l d e d and faulted} beds o~ T,-~rtiary coal~beatin,~ deposits, S.ogo River basin Former meltwater detia, lower Khorogor Riv~.r va}le ~.,
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Strt~c~u~al ,,',eat,u~(-~s in P~sleozoic be(:!
0
5
10
1 5k.m
Figure 11-60. Glaciotectonic thrust systems at Lake Sevastian and Belugalakh Bay, Tiksi area, eastern Siberia. General location map (left). Box (T) indicates position of detailed map of Tiksi region on right. Based on Grosswald et al. (1992, fig.
1).
Aber and Ber
190 commonly folded, faulted, and sheared into mylonite by glacier action. The middle glacial limit, Kurentsovo Line, is a series of sharp-crested ridges, 20-30 km long, 2.5 to 10 km wide,, ld 20-25 m high. Seismic data lead to the conclusion that ridges are composed of upthrust Cretaceous bedrock. The Kurentsovo Line represents an ice-margin related to the Barents Sea ice sheet. Moving east of the Urals into West Siberia, glaciotectonic structures and landforms are ubiquitous over great expanses, including large thrust systems, megablocks, and ice-shoved ridges (Astakhov, Kaplyanskaya and Tamogradsky 1996). A particularly good example is demonstrated in fiver bluff exposures and borehole data along the Ob River in the Atlym vicinity (fig. 11-59). Glaciotectonic disturbances involve a thick sequence of poorly consolidated Tertiary strata, which are overlain unconformably by multiple tills and outwash. Thrust zones can be projected 300 m deep, ice-shoved hills stand up to 100 m high above the local datum, and the zone of deformation can be traced across a distance of 25 km. This site is situated beyond the limit of Weichselian glaciation and must be associated with older ice advances. In eastern Siberia, Grosswald et al. (1992) documented series of ice-shoved hills and associated source basins (hill-hole pairs) in the Tiksi vicinity (fig. 11-60). Two glaciotectonic landscapes are located around Lake Sevastian and adjacent to Belugalakh Bay. Ice-shoved ridges consist of giant slices of black Carboniferous shales displaced toward the southwest and thrust over Permian bedrock. Topographic basins mark discrete sources of shale as embayments on the shore of Lake Sevastian. This same area also contains displaced coalbeating Tertiary strata, which have been folded and faulted. It is clear that glaciotectonic deformations across Siberia took place in frozen substratum under cold-based glaciers, as permafrost is generally 100s of meters deep. According to Astakhov, Kaplyanskaya and Tarnogradsky (1996, p. 165):
Frozen sediments, if clayey and~or icy, can readily deform, thus translating basal glacial stress into sliding of the entire glacier/sediment complex along subglacial shear zones. Buried sediments with high ice content may deform plastically under shear stress, just as glaciers behave. Likewise, watersaturated clay may remain unfrozen and plastic at subzero temperatures. This is especially true where clay contains saline pore water, as in West Siberia. Thus clayey and icerich strata may serve as d6collements beneath a cold-based ice sheet, and the dynamic sole of the glacier may shift down to such d6collements with consequent deformation of the substratum (fig. 11-61). Astakhov, Kaplyanskaya and Tarnogradsky (1996) concluded that large compressional
Figure 11-61. Sequential model for ice-push deformation in a permafrozen glacier bed showing velocity profiles. During ice advance, the dynamic sole of the glacier shifts down into the clay strata. Adapted from Astakhov, Kaplyanskaya and Tarnogradsky (1996, fig. 19). Copyright John Wiley and Sons Ltd. Reproduced with permission. disturbances of this type are commonplace across the vast territory of West Siberia. At this point in time, it is premature to attempt comprehensive mapping of glaciotectonic phenomena across the huge expanses of Arctic Russia. Nonetheless, certain common elements emerge from current knowledge about the region. Major glaciotectonic features generally involve thick sequences of weak or unconsolidated Cretaceous, Tertiary or Quaternary strata, the Carboniferous shale of the Tiksi vicinity being a notable exception. Glaciotectonic features are all situated in the outer zone of glaciation connected with ice sheets derived from Barents Sea or Kara Sea centers. Where these ice sheets advanced into regions with soft substrata, they likely became relatively thin with low-slope surface gradients (Tveranger et al. 1999). Under such conditions, whether warm or cold based, the ice sheets were quite mobile and dynamic, which favored glaciotectonism.
Chapter 12 Dynamism of Glaciotectonic Deformation Fundamental cause of glaciotectonism It now is clear that glaciotectonic deformations may take place in a variety of settings--in front of the glacier, beneath the ice margin, or under the center of a thick ice sheet. Deformations may arise during advancing, maximal, or recessional phases of glaciation. All manner of material from crystalline bedrock to loose Quaternary sediment in both frozen and thawed conditions may be affected. Topographic settings vary from rugged mountains to continental shelves. Ice-shoved material was usually removed from a basin and piled into a hill of some kind. However, the opposite is also known, where an ice-scooped basin was partly filled with material dislocated from surrounding uplands ( R u s z c z y f i s k a - S z e n a j c h 1978). In short, glaciotectonic phenomena may be expected wherever susceptible crustal strata were overridden or pushed by glaciers or ice sheets. Many factors have been cited as important or necessary conditions for glaciotectonic deformation to take place (see Table 11-1). Of these factors, most are related to local topography, substratum material, ice dynamics, or water. These factors often vary considerably over short distances and times. Only the first--lateral pressure gradient-operates everywhere beneath a glacier, regardless of the local nature of topography, substratum material, or ice movement. The lateral pressure gradient is the fundamental cause of glaciotectonic deformation (Rotnicki 1976; van der Wateren 1985). Glaciotectonic deformation takes place when the stress (= pressure) transferred from the glacier exceeds the strength of the material subjected to the stress. A glacier imposes two kinds of stress on its bed: 1) vertical stress due to static weight of the ice column (= glaciostatic pressure), and 2) drag or shear stress due to movement of the ice over its bed (= glaciodynamic stress). The combination of these leads to what Jaroszewski (1991, 1994) called the static-kinematic conception for glaciotectonism. Hart (1995b) identified four zones beneath an ice sheet based on expected basal shear stress and the likelihood for subglacial deformation (fig. 12-1). Most glaciotectonic deformation takes place in the marginal and equilibrium zones, where shear stress is greatest; less deformation develops in the intermediate and divide zones. The manner by which basal shear stress is transferred from the glacier to the substratum
is conditioned by the ice/substratum interface. mechanisms are possible (fig. 12-2).
Three
• No basal slippage and no transfer of shear stress into the substratum. This typically occurs in frozen-bed conditions, and all glacier movement is accomplished by internal ice deformation. • Basal sliding, again with no transfer of shear stress into the substratum. Basal melt water allows the glacier to slip over a rigid bed. • Subglacial deformation, in which displacement takes place within a deformable bed. Thickness and nature of the deformable bed vary considerably; it may be thawed or frozen and may be underlain by a discrete d6collement. Glaciostatic pressure is given by ice thickness (H in m) times density (0.9 g/cc for ice) divided by 10 (see Table 12-1 for symbols and units)" ~zi =
0.9H/10 = 0.09H in kg/cm 2
(1)
This stress equals ~90 kg/cm 2 for each 1000 m of ice thickness. The shear stress created by ice movement is much less, for most situations only 1-2 kg/cm 2, with maximum values up to 10 kg/cm 2 where ice is frozen to its bed or flows around bedrock knobs (Weertman 1961; van der Wateren 1985). 2
0,2
i ~
a
0
I:o
=
.
.
.
.
.
.
.
deformation
Figure 12-1. Idealized profile across an ice sheet showing glacier thickness, basal shear stress, and base of deformation. Four zones: 1 - marginal (0-20 km), 2 - equilibrium (20-200 km), 3 - intermediate (200-800), and 4 - ice divide (8001000 km). Note large vertical exaggeration and different scales for ice thickness and depth of deformation. Adapted from Hart (1995b, fig. 1).
1
9
2
A
b
Table 12-1. Symbols and units used in analysis of glaciotectonic deformation. H
!ee thick~ss (~ight)in m
T
Thickness of substratum in m
p
Pressure or stress: P~ = iit~static pressure, P~= i~ergranular pressure:, P~,,= hydrostatic pressure
o:
kg/cm ~= - I bar or 1 atmosp~re pressure
~ r m a l stress or: pressur,e., Stress ~ i t used ~ r e is kg/cmL !
• i
-
:.-:
:.-:
::
;
;;
.
.
.
.
or to ice dMde; parallel to slope of ice sudace
o.
~ress compor~nt oriented perpendicular (:~rma!) to a plane
S ~ a r stress developed! par:al'!el to a surface of displacement; meas~ed in kglcm~', S ~ a r stress due to ice mo~,~ment over substratum (= glacioT~.,~ dynamic stress) . . . . . . . .
%
Co~sive stre~th of a material (= co~sion}
0
Ar~le of plane (fault) re!ati~ to: the maximum ~rmat stress
t~ .
The lateral pressure gradient can be calculated easily for situations where ice thickness is known (fig. 12-3). Part of the glaciostatic load on the substratum at any point is transferred to a horizontal stress. The relationship between vertical (Crzi)and horizontal (Crx)stresses is given by Poisson's
.
~ r m a l stress in ~rizontai direction; trans~rse to ice margin
o.
.
distance inside the margin. This inequality of ice loading generates a lateral pressure gradient regardless of the direction or rate of ice movement.
ratio (v):
G ia:ciotectonic stress operatic~ ~rizontaIly ,at ice/substratum o~ ,,co~act, Combination of ~ateral pressure gradient (ZAer,)+ :t~.,
.
and Ber
r
~ r m a l stress in ~rtica! direction; stress d ~ to ice badir@ (=: glaciostatic pressure, ~,~), or weight of o~rburden strata (cr~,)
. . . . . . . . .
.
e
A~te of inter~l friction; ~30 ~ for ~ s t rocks .
.
.
.
.
.
.
.
.
.
.
.
.
(~x = ~
((~zi)
For soft sediments and granular materials, Poisson's ratio is close to 0.2, so this value will be used for further calculations (van der Wateren 1985). Thus: (Yx= 0"25~zi = 0.0225H
AH~.~= H~ -H,~
v
Poisson's ratio; describes t ~ ratb of t~rizontal bu~i~ to vertical s~rtening for vertically stressed rock. E×perimentaf val~ ~O.2 for most sedi~,~ary rocks
The vertical load of ice creates a glaciostatic pressure on the substratum, which is irrespective of ice movement. Assuming constant ice thickness over a level surface, the glaciostatic pressure would be equal and uniform in all directions. However, ice thickness is not constant. Ice thickness varies, particularly near the margin, where thickness increases rapidly from zero at the margin to several 100 m a short
(3)
Because the vertical load varies with ice thickness, so the horizontal stress at the glacier sole varies from point to point. The lateral pressure gradient between two points results from the difference (A) in horizontal stresses: m(~xl/2 = (Yxl - (~x2 -"
C # e ~ e in vat~ between points of obserwation, for exampme:
A
(2)
1-v
0"0225(H1- H2)
(4)
For example, consider positions 7 and 8 on the diagram (fig. 12-3) to calculate the horizontal stress difference between two points (Table 12-2). A(~xT/8= (~x7 - Crx8= 13.5 - 10.1 = 3.4 kg/cm 2, or A(~xT/8= .0225(H 7 - Hs)= .0225(150)= 3.4 kg/cm 2
(5) (6)
The horizontal stress differences are cumulative over a distance. In other words, the stress difference over a given interval is passed on and added to the stress difference of the next interval, such that the total horizontal stress difference is:
Figure 12-2. Schematic illustrations for three mechanisms of basal movement of a glacier over the substratum. A -frozen bed with no basal sliding (U F), B - thawed bed with basal sliding (U s), C - subglacial deformation (U D ), which may be either thawed or frozen. Based on Boulton (1996, fig. 1).
FROZEN BED
UNFROZEN ROCK BED
ON ON
Dynamism of glaciotectonic deformation
ICE
o,~,~,.~
"
193
.
.
.
.
Figure 12-3. Cross section showing a hypothetical ice margin with surface gradient and subglacial pressure conditions for 1-km intervals. The summation of lateral pressure differences at the ice margin equals 22.5 kg/cm 2, ignoring the drag of ice movement. Adapted from Aber, Croot and Fenton (1989, fig. 10-1).
",,~+~ X
~*,,~
dX m
.~ kg/cm2
-~
lr
"t
I :1 i Y--,AO" x =
AGxl/2
........"r
t +
AO'x2/3
+
AGx;/4
+
...
. . . . . . . . . . . . . . . . . . . . . . .
TT (7)
In the example (fig. 12-3), the total lateral pressure at the ice margin is 22.5 kg/cm 2. This stress is independent of ice velocity and is controlled entirely by differential ice load. Lateral pressure ideally should be cumulative over long distance, from the center of an ice sheet to its margin. This cannot actually happen, though, because neither ice nor subglacial strata are ideal materials without cohesion or internal friction. Therefore, the distance over which lateral pressure may build is probably limited to a few 10s of km or less. The lateral pressure gradient is greatest near the ice margin, where surface gradient is steepest. This explains why most glaciotectonic features are created at or within a few km of ice margins. Maximum cumulative lateral pressure is probably on the order of 25 kg/cm 2, as it would be rare for ice thickness to change by much more than 1000 m over a distance of 10 kin. For ice sheet interiors, thickness changes over distance are naturally much smaller, but are still present transverse to ice divides and generate small lateral pressure gradients. A thickness change of only 100 m over a distance of 10 km could produce a cumulative lateral pressure of >2 kg/cm 2. Although smaller in magnitude, interior pressure gradients may also produce glaciotectonic disturbances in appropriate substratum materials.
r
"
Table 12-2. Stress factors for idealized glacierprofile, points 7 and 8 (see fig. 12-3). Point
H
7
600
54
:8
450
40.5
.........
....
13.5 10.1
substratum in many situations. To this lateral pressure, the glaciodynamic stress ('Cice)caused by ice movement may be added, assuming ice movement is usually in the same direction as the lateral pressure gradient. Thus, the total glaciotectonic pressure ((yg,) imposed horizontally on the substratum is given by: (~'gt : ZA(~x --I-1:ice
(8)
Given maximum lateral pressure and shear stress values of about 25 kg/cm 2 and 10 kg/cm 2 respectively, this means that maximum glaciotectonic stress ((~gt) is '~35 kg/cm 2. This maximum pressure would be realized in only a few restricted situations, for example where a steep ice front advanced upslope, perhaps over permafrost or during a surge. Many glaciotectonic features were not created under such unusual conditions; lower horizontal stress of
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Figure 13-10. Generalized geological map of the Appalachian thrust belt in northeastern Alabama. Symbols: 1 Lower and Middle Cambrian, 2 - Upper Cambian-Lower Ordovician carbonate (stiff layer), 3 - Middle Ordovician to Lower Mississippian, and 4 - Upper Mississippian-Pennsylvanian (Carboniferous), A T Z - Anniston transverse zone. Across-strike sections (A and B) shown in Fig. 13-11; alongstrike section (C) given in Fig. 13-12. Adapted from Thomas (1990, fig. 4).
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